Induced vs triggered seismicity

The presence of critically stressed faults is thus of great importance for fault reactivation and the occurrence of induced seismicity. It is also one of the challenges, as very small stress changes can generate relatively large seismic events. In the debate on induced seismicity, such relatively large events have also been called ‘triggered’ events (not to be confused with events resulting from dynamic triggering by other earthquakes), whereas events requiring a larger stress change have been called ‘induced’ events. However, it is often unclear what the stress changes were at hypocentre depth, and how the total energy change of a project is related to the released energy. Moreover, a small stress change can advance the occurrence of an earthquake by hundreds to thousands of years. In this article no distinction is made between triggered and induced; from here all the events linked to a project are termed ‘induced’.

Stress changes: pore pressure increase

Pore fluid pressure changes are the most ubiquitous stress perturbation in geothermal systems; pressure changes can occur e.g. due to production and (re)injection of geothermal fluid. An increase in pressure on a fault due to e.g. fluid injection (Fig. 3A) reduces the effective normal stress on the fault and brings the fault closer to failure (Fig. 3D). The elevated pressure diffuses away from the injection well as a function of time, raising pressures and inducing seismic events further away from the source (e.g. Shapiro & Dinske, Reference Shapiro and Dinske2009). Even when injection is terminated, diffusion still causes the pressure to increase at some locations, which may lead to post-injection seismicity. In fact, some of the larger events in Enhanced Geothermal Systems (EGS) have occurred after stimulation was stopped, e.g. at Soultz-sous-Forêts and Basel. When injection occurs in a relatively impermeable but fractured rock mass, the fractures dominate fluid flow and diffusion, whereas for porous rocks the matrix is more important in controlling the fluid flow and diffusion. The pressure distribution thus depends on the rock type targeted and the hydrogeology, as well as on the operational parameters.

Fig. 3.

Simplified examples of the dominant mechanisms causing stress changes in geothermal systems. Examples are given for a normal faulting regime (σ _h < σ _H < σ _v). (A) Schematic 2D illustration of pore pressure increases in a fault and diffusion of pressure along the fault. (B) Schematic 2D illustration of poroelastic stressing in and around a spherical pressurised volume. (C) Schematic 2D illustration of poroelastic stressing in and around a spherical cooled volume. Case 1: stresses on a fault within the pressurised or cooled volume. Case 2: stresses on a fault just outside the pressurised or cooled volume. (D–F) Mohr circle diagram showing the stress changes due to scenarios A–C.

Stress changes: poroelastic stressing

A change in pore pressure also causes a volumetric strain (volume change) of the rock mass experiencing a pressure change. This volume change causes a change in stress within the rock mass itself controlled by pore pressure changes, and in the surrounding rock formations, i.e. poroelastic stressing (Fig. 3B). The magnitude of poroelastic stress changes depends on the pressure change, the elastic properties of the rock mass and the geometry of the rock mass experiencing a pressure (e.g. Segall & Fitzgerald, Reference Segall and Fitzgerald1998; Soltanzadeh & Hawkes, Reference Soltanzadeh and Hawkes2008, Reference Soltanzadeh and Hawkes2009). Figure 3B shows an example for a sphere experiencing an increase in pressure. Inside the volume (case 1) the total stresses become more compressive due to the presence of surrounding rock where the hydraulic pressure is still undisturbed. However, due to the increase in pressure the effective stresses become smaller, and the net effect is a more critical stress state (Fig. 3E). To the sides of the expanding (case 2) volume the horizontal stress becomes more compressive, but the vertical stress decreases, leading to stabilisation in the assumed normal faulting regime. The poroelastic stress-changes outside of the pressurised volume are smaller than inside the volume, and decay rapidly with distance. Poroelastic stressing is expected to play a role both in geothermal systems where pressure is decreased (e.g. producing geothermal fields) or increased (e.g. stimulation in an EGS). Direct pressure effects are expected to dominate near the well, but poroelastic effects reach further at a short timescale. The volume change due to pressure drop can also cause subsidence at the surface, such as observed in many geothermal fields experiencing net extraction, such as the Geysers geothermal field (e.g. Mossop & Segall, Reference Mossop and Segall1997) (USA) and Cerro Prieto (Italy) (Glowacka et al., Reference Glowacka, González and Fabriol1999).

Stress changes: thermoelastic stressing

Analogous to poroelasticity, changes in temperature cause volumetric strain of the rock mass, which leads to stress changes within and around the volume experiencing a temperature change. An example for a spherical cooled volume is shown in Figure 3C. The stress changes are opposite to those resulting from a pressure increase, because cooling leads to a decrease in total stresses. In the normal faulting scenario both faults within and just outside the cooled volume become less stable (Fig. 3D). The effect of temperature changes ΔT may be strong, and is related to pore pressure changes ΔP through (Segall & Fitzgerald, Reference Segall and Fitzgerald1998)

$${{\sigma _{}^{{\rm{thermo}}}} \over {\sigma _{}^{{\rm{poro}}}}} = {{E\lambda {\rm{\Delta }}T} \over {\left( {1 - \upsilon } \right)\alpha {\rm{\Delta }}P}}$$

where E is Young’s modulus, ν is Poisson’s ratio and λ is the linear thermal expansion coefficient. Temperature changes and thermoelastic stressing play an important role in geothermal systems. In geothermal fields the temperature difference between the reservoir rock and reinjected water may be in excess of 200 °C, such as at the Geysers and Larderello (Batini et al., 1985; Martínez‐Garzón et al., 2014). The difference between the reservoir rock and (re)injected fluids in doublets circulating through hot sedimentary aquifers is smaller but still significant (30–70 °C). Over time (10–30 years) these temperature differences may cause a significant fraction of the rock mass in the doublet system (Willems et al., 2017). The thermal stresses in the vicinity of the wellbore can be significant and lead to tensile failure, but the stresses decrease rapidly with distance (Koh et al., Reference Koh, Roshan and Rahman2011). Note that (for tight rocks) thermoelastic and poroelastic stress changes cannot be simply superimposed as they are coupled processes; e.g. cooling will increase the permeability, which influences the pressure distribution.

Other mechanisms of fault reactivation

Beside the three main mechanisms discussed above, other mechanisms may also lead to fault reactivation. These include:

• Mass changes: The addition or removal of mass due to the extraction of fluids or mining causes elastic stress changes at depth which may reactivate (critically stressed) faults (e.g. Segall, Reference Segall1989; Klose, Reference Klose2007; Kang et al., Reference Kang, Zhu and Zhao2019).

• Excavation-induced stresses: Stress changes around an excavated volume (relevant for mining).

• Chemical changes of fault properties: Rather than changing the stress on fault, changes of the fault properties themselves may bring a fault to failure. These changes result from e.g. the introduction of injection fluids or acid into the subsurface, or temperature changes, which can change the friction coefficient (e.g. Kang et al., Reference Kang, Zhu and Zhao2019). Water can weaken silicate-rich fault rocks through stress corrosion (Davis et al., Reference Davis, Nyffenegger and Frohlich1995). Also, dissolution of carbonate can alter the fault properties and cause fault reactivation as proposed for the Molasse Basin (Seithel et al., Reference Seithel, Gaucher, Mueller, Steiner and Kohl2019) or seismicity below impounded reservoirs (Chen & Talwani, Reference Chen and Talwani1998).

• Static and dynamic triggering: As one earthquake is induced, stress is redistributed around the fault area that slipped. The stress changes may promote failure on other faults or fault segments (e.g. Schoenball et al., Reference Schoenball, Baujard, Kohl and Dorbath2012).

• Effects of local geometry: Local heterogeneities may cause the various stress changes to be locally enhanced, e.g. at the interface of layers with strongly different stiffnesses and offset or impermeable faults. For example, subsidence due to large-scale fluid extraction may cause differential strain along faults (Glowacka et al., Reference Glowacka, González and Fabriol1999).

Tectonic setting and geothermal play type

The case studies and geothermal regions included in this review are very different in terms of tectonic regime, heat flow, geology, etc. It is useful to classify the cases according to their geothermal play type; here we follow the classification given in Moeck (Reference Moeck2014), which is based on the geological controls on the thermal regime, heat flow and hydrogeology. Moeck (Reference Moeck2014) distinguishes between conduction-dominated systems and convection-dominated systems, where conduction and convection are indicative of the heat transfer causing the geothermal anomaly. Here we summarise the main characteristics of these systems (Moeck, Reference Moeck2014).

Conduction-dominated systems are found in passive tectonic regions without recent tectonism or volcanism, such as passive continental margins of intracontinental, tectonically quiet areas. Thermal gradients are near-normal, and therefore conduction-dominated systems host mainly low- to medium-temperature geothermal systems. Rock types targeted include mainly sedimentary rocks, and also crystalline basement rocks. In sedimentary basins, heat is transferred mainly through conduction from depth, whereas in crystalline rocks (e.g. granite) heat production is also caused by decay of radioactive elements. Conduction-dominated systems are classified into:

• Intracratonic Basin Type (IBT), which are inactive intracratonic rift basins or passive margin basins (e.g. North German Basin) with a near-normal geothermal gradient. Usually these basins are filled with sediments, which may form permeable aquifers at depths >3 km suitable for the production (and reinjection) of geothermal water. Tighter rock types may require EGS technology (see next subsection).

• Orogenic Belt Type (OBT), which are situated in fold-and-thrust belts and foreland basins (e.g. Molasse Basin). The sediments in the foreland basins are bent and have experienced significant subsidence due to the weight of the nearby mountain belt. Advective heat transport may occur along permeable sedimentary formation from deeper parts to shallower parts further away from the mountain range, causing an elevated geothermal gradient.

• Basement Type (BT), which are related to intrusions or heat-producing element regions (e.g. Cooper Basin). This is an igneous play type consisting of crystalline basement (granite) present below most of the continental area. Natural permeability is often low.

Convection-dominated (CV) systems are high-temperature regions where recent volcanism and/or active plate tectonic processes occur. Examples are magmatic arcs, rift zones, transform fault systems and hot-spot magmatism. Circulation of hot fluids from depth to shallower reservoirs causes an elevated geothermal heat gradient. Convection-dominated systems include:

• Volcanic Field Type (VFT) which are found near magmatic arcs, mid-oceanic ridges, and hot spots (e.g. Icelandic Volcanic Fields, Taupo).

• Plutonic Type (PT) plays, which are related to young orogens and post-orogenic collapse with recent plutonism (e.g. Larderello and the Geysers). Temperatures for both PT and VFT may exceed 300 °C and reservoirs may be vapour- and/or liquid-dominated.

• Extensional Domain Type (EDT) plays, which are situated in metamorphic core complexes, back-arc extension, pull-apart basins and intracontinental rifts (e.g. Upper Rhine Graben, the Great Basin). In contrast to the previous cases, this is a non-magmatic play type, where heat flow predominantly occurs by convection through permeable faults. The crust in these regions is usually thinner, creating a higher geothermal gradient.

Fluid flow and convective heat flow is dominated by matrix porosity (VFT) and/or fractures (EDT, PT).

Figure 4B shows the geothermal play types of the main geothermal areas in Europe. Natural seismicity in the convection-dominated regions is much more ubiquitous than in the conduction-dominated systems, which may make distinguishing natural and induced seismicity difficult (e.g. Brodsky & Lajoie, Reference Brodsky and Lajoie2013).

Geothermal system type

Depending on the tectonic setting and rock type, different types of deep geothermal energy systems exploit the geothermal resources. Here three different types are recognised (following Breede et al., Reference Breede, Dzebisashvili and Falcone2015): hydrothermal systems (HS), petrothermal systems (PS) and hot sedimentary aquifers (HSA) (Fig. 5).

• Hydrothermal systems (HS), also called geothermal fields, are the oldest geothermal systems and are situated in tectonically active, convection-dominated settings (see Fig. 4A). These fields are high-temperature (>200 °C) reservoirs at shallow depth (<∼3 km) and can be classified as vapour-dominated or water-dominated fields. Permeability of the matrix and/or fractures is high enough to allow flow without stimulation. Well-known examples include the Geysers, and Larderello. From the early 1900s, hot water from these fields was used for heating (direct-use) and the first commercial production of electricity was achieved in 1926 at Larderello. The fields can be extensive (>10 km) and contain numerous wells. Initially the fields only produced steam and/or hot water. From the 1970s, reinjection of produced cooled water as a means of disposal and/or pressure maintenance became common practice (Stefansson, Reference Stefansson1997; Kaya et al., Reference Kaya, Zarrouk and O’Sullivan2011). Note that reinjection may not necessarily occur in the producing formation but could also occur shallower or in a different area of the field.

• Petrothermal systems (PS) (here also called EGS) target rocks that have a low permeability and need to be stimulated before fluid flow can be achieved. These include the Hot-Dry Rock (HDR) concept, where wells are connected through hydraulic fracturing which is achieved by injection pressures exceeding the minimum principal stress (e.g. Jung, Reference Jung2013) and Enhanced Geothermal Systems where injection at high pressures (but lower than the minimum principal stress) stimulates the permeability of the natural fracture network between two wells by shearing. Since the 1980s a number of pilot-EGS projects have been conducted, and the first commercial plants are operational (Lu, Reference Lu2018). In this review both initial HDR project and EGS projects are classified as PS. Typically a petrothermal system is exploited by circulation between two wells. Note that there may also be EGSs within geothermal fields, e.g. in the Geysers (PS–HS).

• Hot sedimentary aquifers are porous/permeable sedimentary rock formations where heat flow is dominated by conduction (see Fig. 4A). Mostly these are low–moderate-temperature (30–150 °C) systems in permeable sedimentary aquifers at relatively shallow depth (1–4 km). Typically, water is circulated between two wells (a doublet) at low pressures. Temperatures can be high enough for electricity generation (e.g. in some systems near Munich) but are mostly suitable for direct use in e.g. district heating, space heating and various other applications (e.g. North German Basin, Paris Basin, the Netherlands).

Beside the type of geothermal system, also the depth and reservoir temperature are specified for the case histories included in the review. In addition, the minimum vertical distance of the geothermal system to basement was listed. For HSA and PS this was usually the crystalline basement. For hydrothermal systems where no crystalline basement was present, this could be another rock type, e.g. greywacke (Taupo Volcanic Zone in New Zealand).

Fig. 5.

Schematic illustration of geothermal system types.

Geological parameters

The initial stress is very important for assessing the reactivation potential of faults. However, in situ stress measurements are often not available, and often faults have not been identified. Where available the stress regime was listed, with NF: normal faulting, NF-SS: transtenstional faulting, SS: strike-slip faulting, SS-TF: transpressive faulting, and TF: thrust faulting. For each activity the rock type of the geothermal target formation is listed, as rock type may influence the potential for faulting and controls fluid flow. Also an indication of the average matrix porosity Φ was included. Where no porosity data were available, a standard value of 1% was assumed for granites, 8% for carbonates and 10% for volcaniclastic sequences. To analyse whether there is a relation between tectonic activity and induced seismicity, at each locality the second invariant of the strain tensor (i.e. strain magnitude) was obtained from the Global Strain Rate Model v2 (Kreemer et al., Reference Kreemer, Blewitt and Klein2014) (Fig. 4A).

Operational parameters

Operational parameters are also important for the occurrence of induced seismicity, as larger pressure and temperature changes result in larger stress changes. For each activity we list the start of operations, and the following operational parameters:

• ΔV (m³): net injected volume

• ΔP (MPa): maximum injection pressure at the wellhead

• ΔT (°C): temperature difference between reservoir temperature and (re)injected fluids

Seismicity

If seismic events occurred, the maximum magnitude and the date of occurrence are listed, as well as the magnitude scale that was used. In addition, where available the magnitude of completeness M _c of the local monitoring network was listed.

Conduction-dominated plays and case studies

North German Basin: The North German Basin (Fig. 4B) is a Permian rift basin (IBT) with a thick infill of Mesozoic and Cenozoic sediments (e.g. Scheck-Wenderoth & Lamarche, Reference Scheck-Wenderoth and Lamarche2005). It is a sub-basin of the Central European Basin System. The basin is inactive, and natural seismicity rates are low. The main geothermal target formations are sandstone aquifers from the Lower Jurassic Lias, the Upper Triassic Keuper (which includes the well-known Rhaetian Formation) and the Lower Triassic Bundsandstein sandstones including the Dethlingen, Volpriehausen sandstones (Franz et al., Reference Franz, Barth, Zimmermann, Budach, Nowak and Wolfgramm2018). Starting in 1986 at Waren, five doublets (Neustadt-Glewe, Neubrandenburg, Neuruppin, Sønderborg and Waren) circulate water through the low-temperature (45–100 °C), porous (20–30%) Liassic and Keuper sandstone aquifers at depths of 1–2.5 km (Kabus & Jäntsch, Reference Kabus and Jäntsch1995; Seibt et al., Reference Seibt, Kabus and Hoth2005; Røgen et al, Reference Røgen, Ditlefsen, Vangkilde-Pedersen, Nielsen and Mahler2015; Franz et al., Reference Franz, Barth, Zimmermann, Budach, Nowak and Wolfgramm2018). Apart from an incidental acid stimulation, no stimulation techniques are required and high flow rates can be attained at low injection pressures (<1 MPa). The deeper, tighter Bundsandstein (4.1 km, 150–170 °C) in the west was hydraulically fractured in two geothermal research projects: Genesys Hannover and Horstberg (Orzol et al., Reference Orzol, Jung, Jatho, Tischner and Kehrer2005; Rioseco et al., Reference Rioseco, Löhken, Schellschmidt and Tischner2013; Tischner et al., Reference Tischner, Krug, Pechan, Hesshaus, Jatho, Bischoff and Wonik2013). Similarly, the geothermal potential of older Permian Rotliegend sediments and volcanics was tested by hydraulic fracturing at the Gross Schönebeck research well (e.g. Zimmermann et al., Reference Zimmermann, Moeck and Blöcher2010).

Only very small events occurred during the hydraulic fracturing experiments, with maximum magnitudes of M _w−1 at Gross Schönebeck (Moeck et al., Reference Moeck, Kwiatek and Zimmermann2009; Kwiatek et al., Reference Kwiatek, Bohnhoff, Dresen, Schulze, Schulte, Zimmermann and Huenges2010; Blöcher et al., Reference Blöcher, Cacace, Jacquey, Zang, Heidbach, Hofmann, Kluge and Zimmermann2018), or no events at all (Tischner et al., Reference Tischner, Krug, Pechan, Hesshaus, Jatho, Bischoff and Wonik2013). At the other sites no monitoring was in place, but no events have been reported by the public.

Norwegian–Danish Basin: The Norwegian–Danish Basin (Fig. 4B) is a Permian rift basin which is also part of the Central European Basin System (IBT). The thickness of the sedimentary cover ranges from 1–2 km in the south to 9 km in the north of the basin (Frederiksen et al., Reference Frederiksen, Nielsen and Balling2001). The main geothermal targets are the Upper Triassic Gassum (Rhaetian in Germany) and the Lower Triassic Bundsandstein (Kristensen et al., Reference Kristensen, Hjuler, Frykman, Olivarius, Weibel, Nielsen and Mathiesen2016). Two geothermal plants are operative: Thisted (started in 1984) and Margretheholm (started in 2005) near Copenhagen. Thisted exploits the high-porosity (26%) Gassum sandstone at 1.2 km depth with a temperature of 45 °C (Mahler, Reference Mahler1995; Hjuler et al., Reference Hjuler, Vosgerau, Nielsen, Frykman, Kristensen, Mathiesen, Bidstrup and Nielsen2014; Røgen et al., Reference Røgen, Ditlefsen, Vangkilde-Pedersen, Nielsen and Mahler2015). The Margretheholm plant targets the Bundsandstein at 2.5 km depth (73 °C) (Mahler et al., Reference Mahler, Røgen, Ditlefsen, Nielsen and Vangkilde-Pedersen2013), which directly overlies the granitic basement (Erlström et al., Reference Erlström, Boldreel, Lindstrom, Kristensen, Mathiesen, Andersen, Kamla and Nielsen2018).

No seismic events have been reported for either geothermal plant.

Paris Basin: The Paris Basin is a Permo-Triassic basin in northern France with a sediment thickness up to 3 km in the centre (Fig. 4B). The absence of microseismicity and absence of horizontal motion indicates the basin is tectonically inactive (Cornet & Röckel, Reference Cornet and Röckel2012). The prime geothermal target in the Paris Basin is the ∼200 m thick Middle Jurassic Dogger carbonates (1.5–2 km, 55–85 °C) (Lopez et al., Reference Lopez, Hamm, Le Brun, Schaper, Boissier, Cotiche and Giuglaris2010). The most productive layers in the Dogger strata are permeable oolitic reef deposits and shelf sediments with a net thickness of 15–25 m, with an average porosity of 15%. The fluid flow is generally matrix-controlled, although in some parts fluid flow may locally be fracture-controlled (Rojas et al., Reference Rojas, Giot, Le Nindre, Criaud, Fouillac, Brach and Pauwels1989). The Dogger is underlain by Lower Jurassic shales, Triassic sands and shales, which are underlain by the granitic basement (Fig. 6). Between 1970 and 1985 more than 100 geothermal wells were drilled in the Paris Basin, mainly used for district heating. Of the 55 doublets 34 are still operational, without any thermal decline (Lopez et al., Reference Lopez, Hamm, Le Brun, Schaper, Boissier, Cotiche and Giuglaris2010). Hydraulic stimulation is not necessary, and circulation can occur at low injection pressures (<1 MPa). However, chemical stimulation is sometimes used to improve the injectivity (Ungemach et al., Reference Ungemach, Antics and Papachristou2005).

Fig. 6.

Cross-section of the stratigraphy of the Paris Basin (Dentzer et al., Reference Dentzer, Lopez, Violette and Bruel2016, reprinted with permission from Elsevier).

No (felt) induced events have been reported for any of the doublets in the Paris Basin. No site-specific monitoring is in place.

Polish Lowlands: The Polish Lowlands, or Polish Basin, is another sub-basin of the Central European Basin System, which formed during rifting in the Permian (e.g. Scheck-Wenderoth & Lamarche, Reference Scheck-Wenderoth and Lamarche2005). The basin was filled with >8 km of Mesozoic sediments, with the Polish trough as the deepest part of the basin running NW–SE (e.g. Grad & Polkowski, Reference Grad and Polkowski2016). The main aquifers targeted by geothermal operations in the Polish Lowlands are the porous Lower Jurassic (0.7–3.7 km depth, 25–120 °C) and Lower Cretaceous (0.8–2.8 km depth, 30–85 °C) sandstone formations which have porosities of 5–20% with an average permeability for the Lower Jurassic of 1.1 × 10⁻¹² m² (Kurowska, Reference Kurowska2000; Sowizdzal, Reference Sowizdzal2018). Since 1996 five geothermal plants have become operational that use hot water from these aquifers for district heating (Pyrzyce, Mszczonów, Uniejów, Stargard Szczecinski and Poddebice), as well as several spas (e.g. Meyer and Kozlowski, Reference Meyer and Kozlowski1995; Górecki et al., Reference Górecki, Sowiżdżał, Hajto and Wachowicz-Pyzik2015; Kepinska, Reference Kepinska2015). The systems at Poddebice and Mszczonów (Bujakowski & Wojnarowski, Reference Bujakowski and Wojnarowski2005) only produce the formation water, whereas the others circulate fluids between the wells.

No (felt) induced events have been reported for any of the system in the Polish Basin. No local monitoring stations are present.

Molasse Basin: The Molasse Basin is a foreland basin north of the Alps (OBT), stretching from Switzerland in the west through the south of Germany to Austria in the east (e.g. Reinecker et al., Reference Reinecker, Tingay, Müller and Heidbach2010). The geothermal target formation is the permeable, faulted, karstified Upper Jurassic Malm limestones which underlie the Tertiary Molasse sediments. Near Munich the Malm lies at a depth of 3 km, with a thickness of over 600 m, with temperatures in the range 85–130 °C (e.g. Seithel et al., Reference Seithel, Gaucher, Mueller, Steiner and Kohl2019). Here, the Malm directly overlies the Variscan basement (Fig. 7). Further to the west at Sankt Gallen in Switzerland, Middle Jurassic and Triassic and Permian sediments are found below the Malm (Moeck et al., Reference Moeck, Bloch, Graf, Heuberger, Kuhn, Naef and Wolfgramm2015). So far, a total of 26 geothermal doublets are operational in the Molasse Basin, 16 of which are situated around Munich (Agemar et al., Reference Agemar, Weber and Schulz2014; Baumann et al., Reference Baumann, Bartels, Lafogler and Wenderoth2017; Seithel et al., Reference Seithel, Gaucher, Mueller, Steiner and Kohl2019), and 7 of which are located in the east of the basin near the German–Austrian border (e.g. Goldbrunner, Reference Goldbrunner1999; Karytsas et al., Reference Karytsas, Mendrinos and Goldbrunner2009). The geothermal wells often target faults because of the enhanced permeability. Most of the doublets are used for district heating, and three systems (Dürrnhaar, Kirchstockach and Sauerlach) to the south of Munich generate electricity (Wanner et al., Reference Wanner, Eichinger, Jahrfeld and Diamond2017; Seithel et al., Reference Seithel, Gaucher, Mueller, Steiner and Kohl2019). Circulation occurs at low (1–2 MPa) pressure (e.g. Wanner et al., Reference Wanner, Eichinger, Jahrfeld and Diamond2017), and hydraulic stimulation is not required.

Fig. 7.

Overview of geothermal systems in the Munich area in the Molasse Basin. (A) Overview map showing Munich, the location of geothermal projects, and local fault structures. (B) Cross-section of the Mollasse Basin (Seithel et al., Reference Seithel, Gaucher, Mueller, Steiner and Kohl2019, reprinted with permission from Elsevier).

Three geothermal sites in the Molasse Basin are associated with induced events with M > 2.0 (Fig. 7). The most well-known example is Sankt Gallen, where a M _L 3.5 event was recorded during well control operations following influx of gas (Kraft et al., Reference Kraft, Wiemer, Deichmann, Diehl, Edwards, Guilhem and Woessner2013; Obermann et al., Reference Obermann, Kraft, Larose and Wiemer2015; Diehl et al., Reference Diehl, Kraft, Kissling and Wiemer2017). Other felt events of M _L 2.4 and M _L 2.1 were recorded during circulation in doublets at Unterhaching (Megies & Wassermann, Reference Megies and Wassermann2014) and Poing (LIAG, 2018; Seithel et al., Reference Seithel, Gaucher, Mueller, Steiner and Kohl2019), respectively after a few months and after 5 years of circulation. In both cases the events occurred in the top of the crystalline basement underlying the Malm. A number of other sites were associated with M < 2.0 events: Dürrnhaar (M _L 1.3), Sauerlach (M _L 1.2), Kirchstockach (M _L 0.8), Oberhaching (M _L 0.5), Taufkirchen (M _L 0.3), Pullach, M _L 0.4) (Seithel et al., Reference Seithel, Gaucher, Mueller, Steiner and Kohl2019) and other systems around Munich have not shown any seismicity, even though the monitoring system captures all events with M > 1 .0–1.5 (and locally even smaller events). Also no events have been reported for the doublets in the eastern part of the basin.

Cooper Basin: The Cooper Basin is located in a stable region of the northeast of Australia. The geothermal target formation is a large naturally fractured granitic body covered by 3.6 km of sediments. Heat is generated within the granite, causing high temperature gradients in the area. Two EGS doublets are present at the Habanero geothermal field (Baisch et al., Reference Baisch, Weidler, Voros, Wyborn and Graaf2006, Reference Baisch, Vörös, Weidler and Wyborn2009), and one EGS well in the Jolokia field (Baisch et al., Reference Baisch, Rothert, Stang, Vörös, Koch and McMahon2015).

More than 60,000 events (M _L 1.6 to 3.7) were detected during three hydraulic stimulations: two in Habanero 1 and one in Habanero 4 (Baisch et al., Reference Baisch, Weidler, Voros, Wyborn and Graaf2006, Reference Baisch, Vörös, Weidler and Wyborn2009, Reference Baisch, Rothert, Stang, Vörös, Koch and McMahon2015). Maximum magnitudes were M _L 3.7, M _L 3.0 and M _L 3.0, and the microseismic cloud revealed a large thrust fault structure that was well-oriented in the regional stress field and was reactivated by the injection (Holl & Barton, Reference Holl and Barton2015). Stimulation of the Jolokia well on the other hand resulted in smaller events (M _L < 1.6); here the well did not intersect a large fault (Baisch et al., Reference Baisch, Rothert, Stang, Vörös, Koch and McMahon2015).

EGS Helsinki: Recently, fractured granitic basement rocks of the Svecofennian province (part of the Fennoscandian shield) were targeted with EGS technology below the city of Helsinki, Finland (Kwiatek et al., Reference Kwiatek, Saarno, Ader, Bluemle, Bohnhoff, Chendorain, Dresen, Heikkinen, Kukkonen and Leary2019). Here, the crystalline basement can be found almost up to the surface. Natural seismicity rates in the south of Finland are low, and in the years prior to the stimulation only small natural earthquakes up to M 1.7 were recorded in the vicinity of the drill site. The stress regime is strike-slip. The well OTN-3 was drilled to 6.1 km deep, where temperatures of 125 °C were measured. The well was stimulated by injecting 18,160 m³ of fresh water, at wellhead pressures of 60–90 MPa, and flow rates of 400–800 L min⁻¹.

Seismicity was recorded during the stimulations (M _c − 1.2) and processed near-real-time (Kwiatek et al., Reference Kwiatek, Saarno, Ader, Bluemle, Bohnhoff, Chendorain, Dresen, Heikkinen, Kukkonen and Leary2019). Depending on the evolution of seismic magnitudes in relation to the injected volume, injection flow rates were adjusted. A total of 6150 events were recorded, with a maximum magnitude M _w of 1.9.

Other: Other EGS and HDR test sites targeting granitic basement rocks or tight sandstones included in the database are Fenton Hill (Phillips et al., Reference Phillips, House and Fehler1997), Fjällbacka (Jupe et al., Reference Jupe, Green and Wallroth1992), Paralana (Albaric et al., Reference Albaric, Oye, Langet, Hasting, Lecomte, Iranpour, Messeiller and Reid2014; Riffault et al., Reference Riffault, Dempsey, Archer, Kelkar and Karra2016) and Rosemanowes (Batchelor, Reference Batchelor1982; Richards et al., Reference Richards, Parker, Green, Jones, Nicholls, Nicol, Randall, Richards, Stewart and Willis-Richards1994; Parker, Reference Parker1999; Evans et al., Reference Evans, Zappone, Kraft, Deichmann and Moia2012); see further Table S2 (in the Supplementary Material available online at https://doi.org/10.1017/njg.2019.6). Maximum magnitudes during stimulations at these sites were respectively M 1.3, M _L 0.2, M _L 2.5 and M 2.0. At Fjällbacka, during a later stimulation, an event was felt but its magnitude was unknown.

Convection-dominated plays and/or case studies

Iceland Volcanic Zones: Iceland is the site of active rifting and volcanism (VFT) and has large geothermal potential. The Icelandic volcanic area runs SW–NE across the island and is seismically active. Production of geothermal energy started in the 1970s. Currently eight high-temperature (230–440 °C) fields are being produced (six of them for electricity), and tens to hundreds of low-temperature fields for district heating (Ragnarsson, Reference Ragnarsson2005; Armannsson, Reference Armannsson2016). Reinjection occurs in five high-temperature and four low-temperature fields (Flóvenz et al., Reference Flóvenz, Ágústsson, Guðnason and Kristjánsdóttir2015). The main geothermal reservoir rocks are basaltic lava flows and hyaloclastics, which may have high porosities (e.g. Eggertsson et al., Reference Eggertsson, Lavallée, Kendrick and Markússon2018).

The Iceland Volcanic Zones are well covered by the seismic network, with a magnitude of completeness of ∼1.0 (Panzera et al., Reference Panzera, Mignan and Vogfjörð2017). Seismicity was observed in most of the fields where reinjection occurred, with maximum magnitudes between M _L 2.0 and 4.0 (Flóvenz et al., Reference Flóvenz, Ágústsson, Guðnason and Kristjánsdóttir2015). The largest event, of M _L 4.0, was observed at the Húsmúli reinjection site in the Hellisheidi field (Juncu et al., Reference Juncu, Árnadóttir, Geirsson, Guðmundsson, Lund, Gunnarsson, Hooper, Hreinsdóttir and Michalczewska2018).

Taupo Volcanic Zone: The Taupo Volcanic Zone (TVZ) is a 150 km long NE–SW zone in the centre of the North Island, New Zealand. The TPV is a zone of active rifting and volcanism associated with subduction at the Kermadec trench in the east (Wilson & Rowland, Reference Wilson and Rowland2016). Currently 16 geothermal fields are being developed in the region, mostly targeting pyroclastic, fractured lavas (andesites, rhyolites), and fractured greywacke basement formations at 1–3 km depth (250–320 °C). Porosities vary from several per cent in the deeper rocks to more than 20% in the shallower rocks (Henrys & Hochstein, Reference Henrys and Hochstein1990; Cant et al., Reference Cant, Siratovich, Cole, Villeneuve and Kennedy2018). Production started in the late 1950s from the Kawerau and Wairakei fields, generating electricity and producing warm water for district and industry heating. Reinjection started from the 1980s.

Local monitoring networks are present near most geothermal fields so that most events with M > 2.0 are detected (Sherburn et al., Reference Sherburn, Bromley, Bannister, Sewell and Bourguignon2015a). Induced seismicity was observed in a number of fields, notably Rotokawa (M 3.3), Kawerau (M 3.2), Mokai (M 3.2), Ngatamariki (M 2.7) and Wairakei (M 2.5) (Sherburn et al., Reference Sherburn, Bromley, Bannister, Sewell and Bourguignon2015a). A clear correlation with reinjection wells could be observed for some of those fields (Sewell et al., Reference Sewell, Cumming, Bardsley, Winick, Quinao, Wallis, Sherburn and Bourguignon2015; Sherburn et al., Reference Sherburn, Sewell, Bourguignon, Cumming, Bannister, Bardsley, Winick, Quinao and Wallis2015b). Several other fields in similar settings (Ohaaki, Ngawha) have not shown seismicity.

The Geysers: The Geysers field in Northern California is the largest producing geothermal field in the world. It lies 80 km from the San Andreas Fault between large active strike-slip faults, in a transtensional stress environment (Garcia et al., Reference Garcia, Hartline, Walters, Wright, Rutqvist, Dobson and Jeanne2016). The Geysers field itself is bounded by two NNW–SSE-striking faults that have been inactive for at least 15,000 years, and is underlain by a granitic pluton. The reservoir is vapour-dominated and situated in faulted metagreywacke (tight, competent, metamorphosed, poorly sorted sandstone). The reservoir can be divided into a normal-temperature reservoir (NTR) (∼1–2.6 km, T < 240 °C), and a high-temperature reservoir (HTR) which underlies the NTR in the northwest Geysers (T > 400 °C) (Garcia et al., Reference Garcia, Hartline, Walters, Wright, Rutqvist, Dobson and Jeanne2016). Geothermal production started in 1960, and reinjection started in the early 1980s as field pressures decreased (Majer et al., Reference Majer, Freeman, Johnson, Jarpe, Nihei, Hartline, Walter and Deniliger2017). Currently 60–80% of the water is reinjected. In the northwest of the Geysers an EGS demonstration project is being developed (Rutqvist et al., Reference Rutqvist, Dobson, Garcia, Hartline, Jeanne, Oldenburg, Vasco and Walters2015; Garcia, et al., Reference Garcia, Hartline, Walters, Wright, Rutqvist, Dobson and Jeanne2016).

Since the early 1980s, more than 750 events have been recorded annually in the Geysers geothermal field.

Since 2003, 20 M > 4.0 events have been recorded, with the maximum, M _w 5.0, occurring on 14 of December 2016 (Majer et al., Reference Majer, Freeman, Johnson, Jarpe, Nihei, Hartline, Walter and Deniliger2017). The events show a strong spatial correlation to the (re)injection wells (Martínez‐Garzón et al., Reference Martínez‐Garzón, Kwiatek, Sone, Bohnhoff, Dresen and Hartline2014) and a temporal correlation to injected volumes (Trugman et al., Reference Trugman, Shearer, Borsa and Fialko2016).

Tuscany–Latium: The Tuscany–Latium geothermal region is situated in the west of Italy. The high heat flow derives from the post-orogenic collapse of the Apennines, which causes extension and magmatism (Batini et al., Reference Batini, Brogi, Lazzarotto, Liotta and Pandeli2003). Recent seismic activity is moderate (M 4.1 in 1970), but in 1724 a large event occurred in the region, with intensity up to VII–VIII (Batini et al., Reference Batini, Console and Luongo1985). The northern geothermal fields (Larderello-Travale, Monte Amiata) are characterised by a shallow reservoir (<1 km, T up to 250 °C) composed of evaporitic rocks and limestones, and a deeper reservoir (2.5–3.5 km, T > 350 °C) composed of fractured metamorphic rocks (Giovanni et al., Reference Giovanni, Guido and Adolfo2005). The southern fields (Torre Alfina, Latera, Cesano) target fractured limestones at 0.6–2.0 km depth (T < 230 °C) (Cavarretta et al., Reference Cavarretta, Gianelli, Scandiffio and Tecce1985; Buonasorte et al., Reference Buonasorte, Cameli, Fiordelisi, Parotto and Perticone1995; Evans et al., Reference Evans, Zappone, Kraft, Deichmann and Moia2012). Larderello-Travale is the oldest steam-producing geothermal field, with the first commercial powerplant in 1926.

Seismicity in the fields clustered around (re)injection wells and reached moderate magnitudes up to M _L 3.2 at the Larderello–Travale field (Batini et al., Reference Batini, Console and Luongo1985). A M _w 4.5 (M _L 3.9) occurred in April 2000 near the Piancastagnaio field at Monte Amiata (Mazzoldi et al., Reference Mazzoldi, Borgia, Ripepe, Marchetti, Ulivieri, della Schiava and Allocca2015; Braun et al., Reference Braun, Cesca, Kühn, Martirosian-Janssen and Dahm2018), but the relationship with geothermal operations was difficult to identify due to high natural seismicity in the region and poor available seismic data. Several injection tests were performed, but not all tests resulted in significant seismicity. At Latera several injection tests only induced magnitudes up to 0.8 (Moia, Reference Moia2008).

Salton Sea: The Salton Sea geothermal field (Southern California) is located at the transition from the diverging East Pacific Rise to the strike-slip San Andreas Fault system (McGuire et al., Reference McGuire, Lohman, Catchings, Rymer and Goldman2015). The field is situated in a seismically active transtensional region between strike-slip faults. It is a water-dominated geothermal field located in sandstone, siltstone and shale, which are slightly altered at shallow levels (1–2 km and <280 °C) and more altered and cemented at deeper levels where temperatures exceeded 300 °C. Porosity in the shallow formations is 10–30% but decreases strongly with depth and alteration. Deeper rocks are, however, extensively fractured, which provides permeability. Production of steam started in 1986, and was increased to 5–10 million m³ per month from the 1990s, of which 80% was reinjected. With increasing scale of operations the seismicity rates increased. From 1985 to 2011 over 12,000 earthquakes with M > 1.7 were recorded, mostly between 3 and 7 km depth (Brodsky & Lajoie, Reference Brodsky and Lajoie2013). The largest observed event was a M 5.1. which was part of a large earthquake swarm with multiple M > 4 events. However, the correlation with operational parameters was less straightforward than e.g. at the Geysers; an initial correlation with production rates was observed, but also tectonically driven earthquake swarms occur in and near the field (Brodsky et al., 2013; Trugman et al., Reference Trugman, Shearer, Borsa and Fialko2016).

Great Basin: The Great Basin is a large basin in the west of the USA, characterised by active dextral strike-slip faulting and extensions, i.e. a transtensional stress regime. Geothermal systems in the basin are found near active zones of intersecting or overlapping faults, such as step-over fault systems, which provide permeable pathways for hot fluids, but rarely near mature fault zones, as these may have lower permeability (Faulds et al., Reference Faulds, Coolbaugh, Benoit, Oppliger, Perkins, Moeck and Drakos2010). The main geothermal target formations are metavolcanics and metamorphic (siliceous metamudstone) rocks at 150–250 °C (e.g. Lutz et al., Reference Lutz, Hickman, Davatzes, Zemach, Drakos and Robertson-Tait2010). In a few cases at the western margin, heat flow is of magmatic origin (e.g. at Coso field, T = 285 °C). Here the main reservoir rock is (grano)diorite (Davatzes & Hickman, Reference Davatzes and Hickman2010). Over 400 geothermal systems are found in the Great Basin, 27 of which have temperatures high enough for generation of electricity.

For most systems no induced seismicity is reported, but local monitoring systems were only placed near a couple of geothermal fields. Several small events were measured during EGS stimulation at Desert Peak with M _L up to 1.7 (Benato et al., Reference Benato, Hickman, Davatzes, Taron, Spielman, Elsworth, Majer and Boyle2016) and production and reinjection in Brady’s field with M _L up to 2.4 (Cardiff et al., Reference Cardiff, Lim, Patterson, Akerley, Spielman, Lopeman, Walsh, Singh, Foxall and Wang2018). At Brady, no seismic events were observed when pumping rates were high, but events were temporally correlated to times when production was halted. Larger magnitudes (M < 4.4) occurred in the high-temperature field Coso in the western margin of the basin (Trugman et al., Reference Trugman, Shearer, Borsa and Fialko2016).

Pannonian Basin: The Pannonian Basin (Fig. 4B) is a back-arc basin that started to form in the early Miocene through rifting and volcanism, and covers Hungary as well as parts of Croatia, Slovakia, Slovenia and Romania (Horváth et al., Reference Horváth, Musitz, Balázs, Végh, Uhrin, Nádor, Koroknai, Pap, Tóth and Wórum2015). Present-day inversion causes transpressive stresses, and seismicity occurs in the basin with magnitudes up to M 6.0 (Tóth et al., Reference Tóth, Mónus, Bus, Györi, Husebye and Christova2008). The prime geothermal target is the 100–300 m thick base of permeable (10⁻¹³–10⁻¹² m²) Ujfalu Formation (1–2 km, 30–100 °C), which consists of high-porosity (20–30%) unconsolidated sediments (sands) overlying an aquitard formation (Horváth et al., Reference Horváth, Musitz, Balázs, Végh, Uhrin, Nádor, Koroknai, Pap, Tóth and Wórum2015). A second target is karstified Mesozoic carbonates at depths <2 km. There are over 350 sites producing thermal water mostly from the Ujfalu Formation, of which >150 are used for space heating. Most geothermal systems such as for example Szentes (Szanyi & Kovács, Reference Szanyi and Kovács2010; Bálint & Szanyi, Reference Bálint and Szanyi2015) consist of multiple wells, and only produce water, which over time leads to pressure decreases in the geothermal aquifer. Reinjection occurs in only 20 wells, for example in the Hódmezővásárhely geothermal system (Szanyi and Kovács, Reference Szanyi and Kovács2010) and in the Orosháza-Gyopárosfürdő geothermal system (Szita and Vitai, Reference Szita and Vitai2013).

No seismicity has been reported for any of the geothermal sites. No local networks are in place; the regional network has a detection threshold of ∼M 2.0 (Tóth et al., Reference Tóth, Mónus, Bus, Györi, Husebye and Christova2008).

Pohang: The Pohang EGS site is located in the Pohang Basin in the southeast of the Korean Peninsula and targets a granodioritic intrusion found from 2.4 km depth (Kim et al., Reference Kim, Ree, Kim, Kim, Kang and Seo2018a). The Pohang Basin is a Miocene back-arc basin which is currently subject to a compressive stress regime, and although natural seismicity in the vicinity of the EGS is low, earthquakes with M _w 5.8 have occurred at 40 km from the EGS site (Kim et al., Reference Kim, Min, Kim, Choi, Yoon, Yoon, Yoon, Lee and Song2018b). Two wells (PX-1 and PX-2) were drilled to a depth of 4.2 km with a horizontal distance of 0.6 km between them. The temperature at this depth was 140 °C.

Between 29 January 2016 and 18 September 2017 five hydraulic stimulations were performed, three of them in PX-2 (the first, third and fifth) and two of them in PX-1 (the second and fourth) (Kim et al., Reference Kim, Ree, Kim, Kim, Kang and Seo2018a). The maximum wellhead pressures were reached during the first two stimulations and were 89.2 MPa (PX-2) and 27.7 MPa (PX-1) (Ge et al., Reference Ge, Giardini, Ellsworth, Shimamoto and Townend2019). A total volume of 7135 m³ was injected into PX-2 and 5663 m³ into PX-1, of which 2989 m³ and 3968 m³ was flowed back, so that the net injected volume in both wells was ∼6000 m³. Injection tests indicated that a hydraulic barrier was present between PX-1 and PX-2.

During drilling of PX-2 a fault zone was encountered at 3.8 km depth. This resulted in mud losses, and seismicity of up to M _w 0.9 (Ge et al., Reference Ge, Giardini, Ellsworth, Shimamoto and Townend2019). Seismicity with magnitudes up to an M _w 3.2 in April 2017 was recorded during the various stimulations in PX-1 and PX-2. Injection in PX-2 had reactivated the fault zone which penetrated by the well at 3.8 km depth. The third and last stimulations of PX-2 in September 2017 produced only small-magnitude seismicity up to M _w 2.0 on this fault. However, seismicity restarted in mid-November 2017, with some foreshocks increasing in magnitude from M _w 1.6 to M _w 2.7. Subsequently, on 15 November 2017 the M_w 5.5 mainshock occurred at 4.3 km depth, just below the location of the foreshocks and seismicity previously induced by injection in PX-2 (Grigoli et al., Reference Grigoli, Cesca, Rinaldi, Manconi, López-Comino, Clinton, Westaway, Cauzzi, Dahm and Wiemer2018; Kim et al., Reference Kim, Ree, Kim, Kim, Kang and Seo2018a; Ge et al., Reference Ge, Giardini, Ellsworth, Shimamoto and Townend2019). A first-order hydromechanical modelling approach suggests that the pressure increase at the hypocentre depth was only 0.07 MPa, indicating the fault was close to failure (Ge et al., Reference Ge, Giardini, Ellsworth, Shimamoto and Townend2019).

Upper Rhine Graben: The Upper Rhine Graben is a 300 km long NNE–SSW-trending active rift system on the border of France and Germany. The graben is seismically active, with frequent small events and occasional moderate-sized events, and some historic events of M ∼7.0 (Grimmer et al., Reference Grimmer, Ritter, Eisbacher and Fielitz2017). Five medium- to high-temperature geothermal power plants are currently operational in the graben (Vidal & Genter, Reference Vidal and Genter2018). The wells at Basel (terminated) and Soultz-sous-Forêts target the fractured granitic basement at 3.5–5 km depth with temperatures of 160–200 °C (e.g. Cornet et al., Reference Cornet, Helm, Poitrenaud, Etchecopar and Talebi1997; Baria et al., Reference Baria, Baumgartner, Gerard, Jung and Garnish1999; Häring et al., Reference Häring, Schanz, Ladner and Dyer2008; Dorbath et al., Reference Dorbath, Cuenot, Genter and Frogneux2009; Deichmann et al., Reference Deichmann, Kraft and Evans2014). The geothermal systems at Insheim, Landau and Ritterhofen were developed using a multi-reservoir concept, where the interface of the granitic basement and the overlying tight, fractured sediments are both targeted (e.g. Schindler et al., Reference Schindler, Baumgärtner, Gandy, Hauffe, Hettkamp, Menzel and Wahl2010; Maurer et al., Reference Maurer, Cuenot, Gaucher, Grunberg, Vergne, Wodling and Schmittbuhl2015; Baujard et al., Reference Baujard, Genter, Dalmais, Maurer, Hehn, Rosillette, Vidal and Schmittbuhl2017; Küperkoch et al., Reference Küperkoch, Olbert and Meier2018; Vidal and Genter, Reference Vidal and Genter2018). At Bruchsal, only the tight, fractured Triassic–Carboniferous sandstone sediments were targeted at a depth of 2.5–3.5 km at a temperature of 100–150 °C (Meixner et al., Reference Meixner, Schill, Grimmer, Gaucher, Kohl and Klingler2016). Hydraulic stimulation of the pre-existing fracture network was required except at Bruchsal, and typically volumes between 10,000 and 50,000 m³ are injected at wellhead pressures between 15 and 30 MPa, except at Rittershofen where pressures were lower (Baujard et al., Reference Baujard, Genter, Dalmais, Maurer, Hehn, Rosillette, Vidal and Schmittbuhl2017).

Seismicity was monitored in most geothermal systems. At a number of sites felt events were induced, with the largest event being a M _L 3.4 related to the Basel EGS stimulations (e.g. Häring et al., Reference Häring, Schanz, Ladner and Dyer2008; Deichmann et al., Reference Deichmann, Kraft and Evans2014). At other locations seismicity was smaller, but still felt by the local population, e.g. a maximum M _L 2.9 at Soultz-sous-Forêts after stimulation (e.g. Cornet et al., Reference Cornet, Helm, Poitrenaud, Etchecopar and Talebi1997; Charléty et al., Reference Charléty, Cuenot, Dorbath, Dorbath, Haessler and Frogneux2007; Dorbath et al., Reference Dorbath, Cuenot, Genter and Frogneux2009), a M _L 2.7 at Landau during circulation (Bönnemann et al., Reference Bönnemann, Schmidt, Ritter, Gestermann, Plenefish and Wegler2010) and a M _L 2.4 at Insheim (Küperkoch et al., Reference Küperkoch, Olbert and Meier2018). Events occurred both during stimulation and circulation, and were located predominantly in the granitic basement (Pilger et al., Reference Pilger, Eulenfeld, Brüstle, Wegler, Plenefisch and Schmidt2017; Küperkoch et al., Reference Küperkoch, Olbert and Meier2018). However, felt events did not always occur; at Rittershoffen, close to Soultz-sous-Forêts, the maximum magnitude did not exceed M_L 1.6 (Baujard et al., Reference Baujard, Genter, Dalmais, Maurer, Hehn, Rosillette, Vidal and Schmittbuhl2017; Lengline et al., Reference Lengline, Boubacar and Schmittbuhl2017). At Bruchsal also no seismicity was recorded during circulation, even though a local monitoring network was present with a magnitude of completeness of M _L 0.7 (Gaucher, Reference Gaucher2016).

Japan: Japan is a tectonically active country with active volcanism and has large geothermal potential, with the first commercial geothermal powerplant opening in 1967 at Matsukawa. Currently there are 17 geothermal powerplants operational, and after a lack of interest in geothermal energy in the early 2000s, many more geothermal plants are planned after the Tohoku earthquake and nuclear disaster in 2011 (Tosha et al., Reference Tosha, Nishikawa, Shimada and Oishi2016). Geothermal fields mainly target Quaternary volcanoes and are located in the northeast of the country on the islands of Honshu and Hokkaido (e.g. Yanaizu-Nishiyama, Kakkonda), or in the southwest in Kyushu. In the late 1980s and early 1990s two HDR test sites were drilled and stimulated at Hijiori and Ogachi, targeting granodiorite basement in calderas. Fracturing experiments at Hijiori were performed in 1998, 1992 and 1995, by injection of 1960 m³, 51,500 m³ and 2115 m³ respectively (Sasaki & Kaieda, Reference Sasaki and Kaieda2002). Circulation tests were performed in 1995 and 2000–2002 with recovery factors of 39% and 53% (Kaieda, Reference Kaieda2015). At Ogachi two fracturing experiments were conducted in 1991 and 1992, injecting 10,000 m³ and 5500 m³ respectively (Ito, Reference Ito2003; Kaieda et al., Reference Kaieda, Jones, Moriya, Sasaki and Ushijima2005; Kaieda, Reference Kaieda2015).

At Yanaizu–Nishiyama (Okuaizu) geothermal field, seismicity was monitored by a local network (M _c ∼ 1.0) (Asanuma et al., Reference Asanuma, Mitsumori, Adachi, Saeki, Aoyama, Ozeki, Mukuhira and Niitsuma2011; Mitsumori et al., Reference Mitsumori, Asanuma, Adachi, Aoyama, Ozeki and Saeki2012). Several M > 3.0 events occurred in the field, including a M _JMA 4.9 event in November 2009, after 14 years of production and reinjection. Natural seismicity also occurs in the field, but a correlation with injection has been observed (Asanuma et al., Reference Asanuma, Eto, Adachi, Saeki, Aoyama, Ozeki and Häring2014). Events up to M 2.0 were recorded during fracturing at Ogachi (Kaieda et al., Reference Kaieda, Jones, Moriya, Sasaki and Ushijima2005; Kaieda, Reference Kaieda2015), but event sizes at Hijiori did not exceed M 0.6 during the different fracturing tests (Sasaki & Kaieda, Reference Sasaki and Kaieda2002). However, three months after the long-term circulation test in 2000–2002, a M 2.4 event occurred, at the edge of the stimulated reservoir (Kaieda, Reference Kaieda2015).

Other: Other case histories included in the review include geothermal fields in Mexico and El Salvador. The largest event associated with one of these fields is M _L 6.6 which occurred on the Imperial Valley Fault in the Salton Trough (Glowacka & Nava, Reference Glowacka and Nava1996). Large-scale production from the nearby Cerro Prieto field in Mexico caused a reduction in pressure and subsidence over a large area surrounding the field (Glowacka et al., Reference Glowacka, González and Fabriol1999; Trugman et al., Reference Trugman, Borsa and Sandwell2014). The poroelastic stress-changes and differential subsidence may have contributed to the occurrence of the M _L 6.6 event (Glowacka & Nava, Reference Glowacka and Nava1996). Other fields in Mexico in the Mexican Volcanic Belt and Baja California (Prol-Ledesma & Morán-Zenteno, Reference Prol-Ledesma and Morán-Zenteno2019) include Los Humeros (Lermo et al., Reference Lermo, Antayhua, Quintanar and Lorenzo2008; Urban & Lermo, Reference Urban and Lermo2012, Reference Urban and Lermo2013; Flores-Armenta, Reference Flores-Armenta2014), Los Azufres (Iglesias et al., Reference Iglesias, Contreras, Garcia and Dominquez1987; Arellano et al., Reference Arellano, Torres and Barragán2005; Urban & Lermo, Reference Urban and Lermo2012), and Los Tres Virgenes (Verma et al., Reference Verma, Pandarinath, Santoyo, González-Partida, Torres-Alvarado and Tello-Hinojosa2006; Antayhua-Vera et al., Reference Antayhua-Vera, Lermo-Samaniego, Quintanar-Robles and Campos-Enríquez2015; Prol-Ledesma et al., Reference Prol-Ledesma, Arango-Galvan and Torres-Vera2016). At Los Humeros, only a small fraction of the extracted fluid is reinjected, at Los Azufres about 50% is reinjected and at Los Tres Virgenes almost 80% is reinjected. Maximum observed magnitudes in these fields are M _d 2.5, M _d 4.6 and M _d 1.9 respectively. In El Salvador, electricity is generated at two geothermal fields: the Chipilapa–Ahuachapan (Fabriol & Beauce, Reference Fabriol and Beauce1997; Monterrosa & López, Reference Monterrosa and López2010) and Berlín (Rivas et al., Reference Rivas, Castellón and Maravilla2005; Rodríguez & Aníbal, Reference Rodríguez and Aníbal2005; Kwiatek et al., Reference Kwiatek, Bulut, Bohnhoff and Dresen2014) fields. Maximum magnitudes detected were M _d 3.0 and M _w 3.7 (M _L 4.4) respectively. The largest event in Berlín followed a hydraulic stimulation of one of the injection wells in the field (Kwiatek et al., Reference Kwiatek, Bulut, Bohnhoff and Dresen2014). At the Mirvalles field in Costa Rica microseismicity increased over time, with a maximum magnitude of M _L 3.8 in 2003 (Moya & Nietzen, Reference Moya and Nietzen2010; Moya & Taylor, Reference Moya and Taylor2010). At the Palinpinon geothermal field in the Philippines, seismicity up to M 2.4 was observed after starting production and injection in 1983 (Bromley et al., Reference Bromley, Pearson and Rigor1987).

Summary

The occurrence of seismic events with M _max> 2.0 for the different geothermal plays and systems discussed above is summarised in Figure 8. Most seismic events have been observed in the convection-dominated play types, mostly in the VFT and EDT (Fig. 8A). In these settings mostly hydrothermal systems (HS) are found, as well as some petrothermal systems (PS), which both show a relatively high number of M > 2.0 sites (Fig. 8B). Fewer events have been observed in conduction-dominated plays (IBT, OBT and BT). In particular for IBT no events of M > 2.0 have been reported to date. Seismicity was also mostly associated with projects targeting crystalline or metamorphic rocks (Fig. 8c). However, as observed from the review, PS in granite does not always induce M > 2.0 events; e.g. at Jolokia, Rittershoften and Helsinki, magnitudes did not exceed 1.9. Hot sedimentary aquifers (HSA) are only occasionally associated with seismicity of M > 2.0. The three cases of seismicity in HSA shown in Figure 8C are from the Molasse Basin where the underlying basement was reactivated. Operations in sandstone were not associated with a large number of M > 2.0 events. For HSA systems in sandstones, no M > 2.0 seismicity was reported, but for PS in Paralana, and HS in Mexico and USA, seismicity occurred. For the latter two cases magnitudes exceeded M 5.0.

Fig. 8.

Occurrence of induced seismic events for different (A) play types, with VF: volcanic fields, PT: plutonic type, EDT: extensional domain type, IBT: intracratonic basin type, OBT: orogenic belt type, and BT: basement type. (B) Rock types, and (C) system types, with PS: petrothermal systems or EGS, GF: geothermal field, and HSA: hot sedimentary aquifer.

Effect of geological parameters

For the cases presented in this study, felt seismicity in conduction-dominated settings only occurred in projects targeting rocks with low matrix porosity at depths larger than 2.5 km (Fig. 9A). The largest magnitude, of M 3.7, occurred due to stimulation of the Habanero EGS in the granite basement of the Cooper Basin (#2 in Fig. 9A). Also a few of the projects in the fractured carbonate rocks of the Molasse Basin generated seismicity, for example at Sankt Gallen (M _L 3.5, #23), Unterhaching (M _L 2.4, #26) and Poing (M _L 2.1, #21). Proximity and a hydraulic connection to the granitic basement resulted in seismic events in these cases. However, for most of the fractured carbonates of the Molasse Basin no seismicity was reported, even though operations (circulation) are similar. This underlines how induced seismicity can be site-specific. Also, deep and tight sandstones of the North German Basin that were hydraulically fractured did not produce felt events. Felt seismicity was also not observed for the 35 doublets targeting the relatively porous Dogger carbonates in the Paris Basin (#45) which have been operational for >30 years. Shallow porous sandstone aquifers in the North German Basin (#35–39), Norwegian–Danish Basin and in the Netherlands (WNB) have also not been associated with felt seismicity.

Fig. 9.

Average rock matrix porosity against depth and the occurrence of induced seismicity. Cases where induced seismicity occurred are shown by square symbols where the size of the symbol scales with maximum magnitude. Cases where no seismicity was observed or seismic events were low (M < 2.0) are indicated with circles. Colours of the symbols give the rock type. (A) Case histories in conduction-dominated settings, including EGS and hot sedimentary aquifers (HSA). (B) Case histories in convection-dominated settings including EGS and hydrothermal systems (geothermal fields). For numbers see Table S2 (in the Supplementary Material available online at https://doi.org/10.1017/njg.2019.6); geothermal areas PL: Polish Lowlands, URG: Upper Rhine Graben, WNB: West Netherlands Basin.

Felt seismicity was more often observed in geothermal projects in convection-dominated settings in tectonically active areas (Fig. 9B), in agreement with Figure 8A. Here, felt seismicity occurred in petrothermal systems (e.g. EGS projects) targeting low-porosity rocks, as well as in hydrothermal fields. The EGS projects targeted naturally fractured granites with low matrix porosity at depths > 3.0 km, including projects at Soultz-sous-Forêts (M _L 2.9) and Basel (M _L 3.4). The largest event was observed at the EGS near Pohang (M _w 5.5, #46). The geothermal fields were located at shallower depths than the EGS, with maximum depths between 2 and 3 km. The rock types and porosities vary; some rocks such as the ignimbrites in the Taupo Volcanic Zone have a high porosity, whereas metamorphic rocks have lower porosities. Seismicity is common, and often clusters around reinjection wells (e.g. Rotokawa, the Geysers, Larderello, Hellisheiði). The largest event occurred in the Salton Sea geothermal field (M 5.5) which is composed of hydrothermally altered sandstones (#60), and at the Geysers (M 5.0, #7). The occurrence of natural seismicity sometimes makes it difficult to distinguish natural and induced seismicity, as for example at Salton Sea. Although seismicity is common, it is not always of large (M > 4) magnitude.

Figure 10 shows the maximum induced magnitudes against various geological parameters. Most felt induced events occurred in areas which were tectonically active (Fig. 10A), in agreement with Figure 9. All M > 4.0 events occurred in regions with a strain rate of 5 nanostrains a⁻¹ or more. However, no further trend between the tectonic loading rates and maximum magnitudes was observed. The M _w 5.5 at the Pohang EGS (#46) occurred in an area with relatively low tectonic loading rates of ∼5–6 nanostrains a⁻¹, which is more than three orders of magnitude lower than the strain rate magnitude at e.g. the Geysers (M 5.0, #7) and Salton Sea (M 5.1, #60). Projects in tectonically inactive regions (R in Fig. 10A) on the other hand also generate felt events, up to a M _L 3.7 at Habanero (#2). Both at Pohand and Habanero the pre-existing faults were close to failure, indicating that low tectonic loading rates do not mean that the stress is low. Projects in tectonically active regions on the other hand do not always generate seismicity, e.g. in the Pannonian Basin. Other factors such as the lithology and operational parameters are also of influence on seismicity.

Fig. 10.

Geological and tectonic parameters against maximum magnitudes observed at geothermal sites. Colours indicate the rock type of the geothermal target formation, the symbols the system type. The dotted lines indicate the range of possible magnitudes for cases where no seismicity was reported. The upper bound of this range indicated the magnitude of completeness in case local networks were present, or a M 2.0 (approximately the threshold for events to be felt) if no local network was present. (A) Depth to (crystalline) basement (negative values mean below basement). (B) Strain rate magnitude (second invariant of the strain rate tensor) from the Global Strain Rate Model (GSRM) (Kreemer et al., Reference Kreemer, Blewitt and Klein2014). R indicates the geothermal site was situated in a region assumed inactive in the GSRM. (C) Tectonic regime, where NF: normal faulting regime, NF-SS: transtensional regime, SS: strike-slip regime, SS-TF: transpressive regime, TF: thrust faulting regime.

Felt events associated with petrothermal systems (PS) and hot sedimentary aquifers (HSA) showed a clear relationship with proximity to the basement (Fig. 10B). All felt events occurred at sites which targeted depths close to the basement, whereas no felt events were reported for PS and HSA targeting formations >0.5 km away from the basement. A M _L 3.5 occurred ∼0.5 km below the Sankt Gallen well; here a previously undetected fault caused the pressure perturbation to migrate to depth. This shows the importance of faults as hydraulic conduits to depth. For hydrothermal systems, on the other hand, the distance to (crystalline or metamorphic) basement had no apparent influence on the occurrence of felt events. No clear trend was visible for event magnitudes and tectonic regime.

Effect of operational parameters

For PS and HSA the maximum magnitudes of induced seismicity increased with the injection pressure ΔP at the wellhead (Fig. 11A). Note, however, that ΔP is closely linked to the rock type targeted for the geothermal operations. Porous sandstones and permeable carbonates are targeted in HSA, which require relatively low ΔP, generally <5 MPa. Granites targeted by EGS on the other hand require >10 MPa during stimulation. Most felt events occurred at ΔP > 5 MPa, in systems targeting crystalline rock. However, seismicity was also induced at low ΔP (<1.5 MPa) during circulation at the aforementioned HSA systems at Unterhaching and Poing, as well as at Insheim in the Upper Rhine Graben. In all these cases, circulation occurred close to the crystalline basement, and most seismicity was located within the (critically stressed) basement rather than the sedimentary formation (Megies & Wassermann, Reference Megies and Wassermann2014; Küperkoch et al., Reference Küperkoch, Olbert and Meier2018). Both the magnitude of the pressure change, as well as the rock type and in situ stress, are thus important for the occurrence of felt events. For hydrothermal systems the relation with ΔP and induced magnitudes is not obvious. Note that e.g. at the Geysers no injection pressure is required; the reservoir formation is underpressured and water can be injected under gravity drive (here shown as ΔP 0.1 MPa). However, at the reservoir level the pressure change with respect to the initial formation pressure may be larger due to the water column, e.g. 7 MPa which is an excess pressure of ∼3.5 MPa (Martínez‐Garzón et al., Reference Martínez‐Garzón, Kwiatek, Sone, Bohnhoff, Dresen and Hartline2014). For future studies we recommend taking the pressure difference with respect to the initial reservoir pressure, rather than the wellhead pressure (the latter is equal to the former in case of a hydrostatic pressure gradient, but not for underpressured or overpressured reservoirs).

Fig. 11.

Operational parameters and maximum magnitudes of induced seismicity at the sites included in the case study review. The dotted lines indicate the range of possible magnitudes for cases where no seismicity was reported. The upper bound of this range indicated the magnitude of completeness in case local networks were present, or a M 2.0 (approximately the threshold for events to be felt) if no local network was present. (A) Injection pressure at the wellhead, (B) Absolute volume change, (C) Maximum reservoir temperature, (D) Temperature difference between reservoir temperature and injected fluid temperature.

There was no clear correlation of M > 2.0 events with the absolute volume change in the geothermal systems (Fig. 11B). However, most M > 4.0 events occurred in hydrothermal systems in which large volume changes (often depletion) occur. One exception is the M _w 5.5 Pohang event, in which the net injected volume was ∼6000 m³, which was comparable to the volume change of other EGS. Some hot sedimentary aquifers have large volume change but no reported seismicity (dotted lines), further illustrating the lack of a clear correlation between volume change and felt seismicity. For injection-induced seismicity in general, such a correlation was observed (McGarr, Reference McGarr2014; Zang et al., Reference Zang, Oye, Jousset, Deichmann, Gritto, McGarr, Majer and Bruhn2014); perhaps here the range of volume changes is too limited to clearly see that correlation.

In this study, all events with M > 2.0 occurred in systems targeting formations of ∼100 °C and hotter (Fig. 11C and D). Temperature differences are in excess of 80 °C, up to more than 200 °C in hydrothermal systems. The data indicate a trend of increasing maximum magnitudes with increasing temperature (and in most cases associated ΔT). Again, the M _w 5.5 Pohang event is an outlier to that trend, i.e. it occurred at relatively low temperatures compared to the other sites.

Doublet operations in the Netherlands

For geothermal doublets in the Netherlands the pumping pressure applied for the thermal loop (production plus injection) is in the order of 20% of reservoir hydrostatic pressure, due to performance and safety constraints (Van Wees et al., Reference Van Wees, Kronimus, Van Putten, Pluymaekers, Mijnlieff, Van Hooff, Obdam and Kramers2012). The allowed pressure changes at the injector are restricted to 13.5 MPa km⁻¹ (State Supervision of Mines & TNO-AGE, 2013). For a saline water gradient of 10.8 MPa km⁻¹ and a reservoir depth of 2 km this would, for example, give a pressure difference of 5.4 MPa at the reservoir level. For high-porous matrix-dominated reservoirs at a few hundred metres from the well, the pressure changes are likely to be small, in the order of <0.1 MPa (Willems et al., Reference Willems, Nick, Weltje and Bruhn2017). However, temperature changes can be significant; simulations of a doublet in the West Netherlands Basin show how the aquifer temperature around the injection well can decrease by more than 30 °C within an area of several hundreds of metres over 30 years (Willems et al., Reference Willems, Nick, Weltje and Bruhn2017). Such changes can cause considerable thermoelastic stressing in the cooled reservoir-volume surrounding the injection well (see ‘Stress changes: thermoelastic stressing’ above). The thermal stress effect is strongly dependent on thermal cooling (ΔT) and the Young’s modulus (E) of rocks (Equation 1), and may thus be stronger with depth (higher ΔT and more competent rocks). Further research on the effect of thermal stresses on the potential for fault reactivation is recommended: for example on the spatiotemporal distribution of the stress changes and the stress path, and to what extent these stresses may be transmitted to or relaxed by inelastic deformation of over- and underlying clays.

In many of the geothermal regions in the Netherlands the seismic network has a magnitude of completeness of 1.5 (Dost et al., Reference Dost, Goutbeek, Van Eck and Kraaijpoel2012). Locally the detection completeness can be lower, in particular in the north of the Netherlands near the Groningen gas field.

In the following subsections the main geothermal target formations in the Netherlands are described. For information on the nomenclature and lithological descriptions see also www.dino-loket.nl.

Tertiary

Tertiary rocks (Palaeogene, Neogene) occur nearly everywhere in the Netherlands at depths of 0–1300 m, though in the RVG the depth increases locally up to 1.8 km (Fig. 13A). The thickness is mostly up to about 100 m, but in the RVG and WNB it can increase to 200–300 m. The rocks constitute loosely consolidated sandstone, siltstone and claystone of predominantly marine origin. The Tertiary rocks unconformably overlie the Cretaceous Chalk group, and are usually unconformably overlain by the sandstones, siltstones and claystones of the Quaternary Upper North Sea Group. Due to the relatively shallow depth, the temperature of the reservoirs is also relatively low. In the past the reservoirs have not attracted much attention for geothermal exploration, but the first well into the Tertiary (Zevenbergen) was drilled in 2017 to a depth of 0.8 km, reaching a temperature of 30 °C. Production is expected to start in 2019. The loosely consolidated sandstone and shale and clays are very porous and have a low competency, and pressure and stress changes are likely low. A hydraulic connection to competent (basement) rocks at depth is not present. Operations in similar lithologies such as in the Pannonian Basin have not generated any seismicity, and the seismogenic potential of this formation is low.

Fig. 13.

Depth maps of the main geothermal target formations in the Netherlands; doublets targeting these formations are indicated by black dots. Depth maps of the aquifers are from ThermoGIS v2.1. (A) Tertiary (Brussels Formation), (B) Lower Cretaceous (Rijnland Group), (C) Upper Jurassic (Schieland Group, Delft and Alblasserdam Members), (D) Lower Triassic, (E) Permian (Rotliegend), (F) Upper Carbonifeorus (Limburg Group, Hunze and Dinkel Subgroups), (G) Lower Carbonfierous (Zeeland Group), (H) cross-section along the AA–AA′ line shown in (E) (www.dino-loket.nl).

Lower Cretaceous and Upper Jurassic sandstones

The Cretaceous and Upper Jurassic sandstones are a prime target of geothermal operations in the Netherlands, and have shown favourable conditions in the WNB (Fig. 13B and C). The WNB is a small transtensional basin containing a number of NW–SE-striking half-grabens that formed during the Jurassic and Early Cretaceous (e.g. Willems, Reference Willems2012).

The doublets drilled to these formations have reached depths of 2–3 km, reaching temperatures between 60 and 90 °C (Table S1, in the Supplementary Material available online at https://doi.org/10.1017/njg.2019.6). Figure 14 shows the stratigraphy of the WNB, including the main target formations which are the Vlieland Sandstone Formation (with the Berkel Sandstone and Rijswijk Members) and Nieuwerkerk Formation (with the Delft and Alblasserdam Members), which are comprised of fluvial sands interbedded with clay layers with average porosities of 20% (DeVault & Jeremiah, Reference DeVault and Jeremiah2002, Willems, Reference Willems2012; Vondrak et al., Reference Vondrak, Van de Weerd, Van Leeuwen and Brautigam2014; Donselaar et al., Reference Donselaar, Groenenberg and Gilding2015). The Upper Jurassic is underlain by the Aalburg Formation of the Altena Group, which consists of calcareous claystones (www.dino-loket.nl). The presence of these thick clay-rich sequences and the depth of the basement (>5.5 km) makes a hydraulic connection to the basement unlikely. The basin is tectonically inactive, and the stress is extensional with the maximum horizontal stress oriented NW–SE. The stress magnitudes are unknown. Several oil and gas fields targeting the Lower Cretaceous and Upper Jurassic formations are present in the area (e.g. Fig. 1B), such as the Berkel Field, De Lier Field and the IJsselmonde Field (www.nlog.nl, de Jager et al., Reference De Jager, Doyle, Grantham and Mabillard1996). These fields, although they are located somewhat shallower than the geothermal doublets, have not led to any induced events over decades of hydrocarbon production, suggesting the stress is not critical. Also the geothermal doublets have been operational from 2006 without inducing (felt) seismicity. In similar geological settings in the North German Basin and Poland no events have been recorded. These observations suggest that geothermal operations in the shallow porous sandstones of the West Netherlands are unlikely to generate seismicity.

Fig. 14.

Stratigraphy of the West Netherlands Basin, from the Late Triassic to the Cenozoic (DeVault & Jeremiah, Reference DeVault and Jeremiah2002). AAPG©2002. Reprinted with permission of the American Association of Petroleum Geologists (AAPG) whose permission is required for further use.

Triassic sandstones

Another geothermal target formation is the Triassic Buntsandstein Group (Fig. 1D), including the Nederweert Sandstone, Volpriehausen Formation, Dethlingen Formation and Hardegsen Formation, which are comprised of a succession of sandstones and silty claystones. The Main Buntsandstein Group is underlain by a 150 m thick Lower Triassic succession of clays and sands, and overlain by evaporites, claystones and carbonates of the Upper Germanic Trias Group. The Buntsandstein Group has favourable conditions for geothermal exploitation in the WNB and possibly in the RVG in the southeast (e.g. Kramers et al. Reference Kramers, Van Wees, Pluymaekers, Kronimus and Boxem2012), with temperatures of 75–130 °C and depths of 2–4.5 km. However, due to a burial anomaly the porosity and permeability are low in the central parts of the WNB and the RVG in the southeast. At the margins of the WNB porosity is good, ranging from 10% at 3 km to 20% at 2 km depth (Bender, Reference Bender2012). Here the Vierpolders doublet (Fig. 13C) has been circulating since the start of 2016. The state of stress is unknown, but no induced seismicity has been observed in relation to geothermal operations, or in relation to oil and gas fields targeting the Buntsandstein Formation in the WNB such as the Brakel Field (www.nlog.nl). In the Brakel Field, hydraulic fracturing was performed in 2008 with injection pressure >25 MPa, but no seismicity was reported and the state of stress was likely not critical. The distance between geothermal operations in the Trias in the WNB and basement is large (>2 km) and a hydraulic connection to basement is thus unlikely. Pressure changes are low, but temperature change may be significant. In similar geological settings in the North German Basin and Poland no events have been recorded. These factors indicate that geothermal operations in the Triassic sandstones of the West Netherlands are unlikely to generate felt seismicity, but it remains important to avoid large faults. Note however that in other areas (CNB, northern Netherlands) gas production from the Buntsandstein Formation has led to felt seismic events, notably in the Roswinkel Field (up to M _L 3.4), and the offshore Q04-A field (up to M _L 2.7) and Q01-B field (up to M _L 2.6) (Fig. 1B). We recommend further research into the response of the Triassic Buntsandstein Formation to geothermally induced stress changes, and potential differences in geological and petrophysical parameters and stress between the WNB and other areas in the Netherlands.

Permian sandstones

The Permian Slochteren sandstone Formation of the Permian Rotliegend has favourable temperature–depth conditions in the CNB and TIJH. Six doublets are currently producing in the provinces of North Holland and Flevoland (Fig. 13E): the three Middenmeer doublets, Koekoekspolder (Veldkamp et al., Reference Veldkamp, Mijnlieff, Bloemsma, Donselaar, Henares, Redjosentono and Weltje2015), Andijk, and Luttelgeest. These doublets target depths between 1.9 and 2.3 km where temperatures are 70–100 °C (Table S1 in the Supplementary Material available online at https://doi.org/10.1017/njg.2019.6). The Slochteren is overlain by anhydrites, carbonates and rock salt, which likely have stable stresses due to viscoelastic relaxation. Here the Upper Carboniferous underburden (Limburg Group) is composed of sands, shales and mudstones (Ruurlo Formation, www.dinoloket.nl). Seismic events have been observed for gas fields producing from Permian sandstones in the north and west of the Netherlands at a depth of 3 km. These include the Groningen gas field with >1000 earthquakes with magnitudes up to M _L 3.6, the Bergermeer field with several M _L > 3 events, the Eleveld field (up to M _L 2.8), and the Annerveen field (up to M _L 2.4). Many of these events occurred after decennia of production, which can be an indication that the state of stress is initially not critical. The formation can be seismogenic, but the stress path and stress changes due to gas depletion are different than those in geothermal systems. We recommend further research into the differences and similarities between the stress changes generated during gas depletion and geothermal operations; in particular the thermoelastic stress which can be large. Seismogenic potential of the doublets is likely low, but care must be taken to stay away from faults (see also e.g. Baisch et al., Reference Baisch, Koch, Stang, Pittens, Drijver and Buik2016; Wiemer et al., Reference Wiemer, Kraft, Trutnevyte and Roth2017). Care must be taken also in case of geothermal activities in or near (former) gas fields, as gas depletion may have altered the state of stress.

Upper Carboniferous

A potential target formation is the Upper Carboniferous sandstone rocks of the Hunze and Dinkel Groups (Fig. 13F). These occur mainly in the east and WNB, whereas in the rest of the Netherlands the Hunze and Dinkel Group are characterised by shales. The sandstones were formed as sheet flood and fluvial channel sandstones. The aquifer is underlain by shales of the Epen Formation, and overlain by sandstones of the Rotliegend. The expected permeabilities are low, with the exception of the eastern part of the Netherlands, coinciding with a shallow depth and therefore relatively low temperature. In the largest part of the WNB, the rocks are buried too deep for conventional geothermal doublet operations without the necessity of stimulation.

Lower Carboniferous

The deepest formation currently targeted in the Netherlands is the fractured and karstified Zeeland Formation of the Carboniferous Limestone Group. In large parts of the Netherlands the Zeeland Formation lies at depth >4 km (Fig. 13F), and its composition in these parts is unknown (Fig. 13G). In the province of Limburg in the southeast of the Netherlands, karstified, fractured and faulted Dinantian limestones of the Zeeland Group are targeted by two doublets: Californië Wijnen Geothermie (CWG) which started in January 2013, and Californië Lipzig Gielen (CLG) which started in the summer of 2017 (Table S2, in the Supplementary Material available online at https://doi.org/10.1017/njg.2019.6; Burghout et al., Reference Burghout, Vorage and Broothaers2019; Vörös and Baisch, Reference Vörös and Baisch2019). The doublets are situated in a stable fault block (Peel and Venlo blocks) north of the RVG and reach a depth of ∼2.5 km (Fig. 13F). Matrix porosity is low, but permeability is provided by the fractures and faults. The production wells of both doublets, and the injection well of CWG cross-cut the Tegelen Fault, which is a large fault bounding the Venlo Block (Vörös & Baisch, Reference Vörös and Baisch2019). Natural seismicity occurs in the region, with eight events with M > 2.2 within 20 km of the doublets. Slip tendency analysis indicates that the state of stress on (some) faults including the Tegelen Fault is close to failure (Vörös & Baisch, Reference Vörös and Baisch2019). At this location the Zeeland Formation is overlain by claystones (Epen Formation), and competent Devonian sandstones/quartzites may be present in the underburden (Reith, Reference Reith2018; Burghout et al., Reference Burghout, Vorage and Broothaers2019).

Even though pressure changes during circulation are small, the fact that circulation occurred through a large fault zone motivated local monitoring of seismicity, which was conducted from 2014 (Baisch & Vörös, Reference Baisch and Vörös2019; Vörös & Baisch, Reference Vörös and Baisch2019). Operations in CWG were suspended in May 2018. After a small (M _L 0.0) event occurred near CLG in August 2018, operations were halted. Within two weeks, nine more events occurred up to M _L 1.7. Since then, operations in both doublets remained suspended because of the uncertainties related to the nature of the seismic events (State Supervision of Mines, 2019). Although the events appear to occur several km below the geothermal reservoir (Spetzler et al., Reference Spetzler, Ruigrok, Dost and Evers2018; Baisch and Vörös, Reference Baisch and Vörös2019), they are temporally correlated to production stops in the CWG doublet. A possible explanation could be the interplay between pressure changes and thermal stresses (Baisch & Vörös, Reference Baisch and Vörös2019; Vörös & Baisch, Reference Vörös and Baisch2019).

Doublet operations in carbonates abroad have occasionally generated seismicity with M > 2.0. Recently, injection tests in a geothermal well targeting the same limestone in Belgium (Balmatt) have generated events up to M _L 2.1 (23 June 2019), clearly lining up on a linear structure close to the injection well signifying that a fault has been reactivated (www.vito.be). Doublets in the Malm carbonates of the Molasse Basin near Munich have in a few instances (2 out of 16) caused induced events (M 2.0–2.4). Most of these larger events occurred in the underburden of the Malm, i.e. in the crystalline basement. This is different from the situation in the Netherlands where other (competent) sediments underlie the Lower Carboniferous Limestone instead of crystalline rocks (though stressed, competent sedimentary formations may also host seismic events). Smaller events have occurred in the vicinity of other doublets in the Malm (Seithel et al., Reference Seithel, Gaucher, Mueller, Steiner and Kohl2019). Fault analysis showed that the occurrence of seismicity is site-specific and depends, among other things, on the structural fabric and state of stress (Seithel et al., Reference Seithel, Gaucher, Mueller, Steiner and Kohl2019).

The observations from the Netherlands and abroad show that (small to moderate) seismic events may thus occur in or around geothermal operations in carbonate formations, and the potential for induced seismicity is larger than for the porous sandstone aquifers. It is important to characterise large faults in the vicinity of the (planned) geothermal system, and their reactivation potential in the stress field.

The review showed that geothermal operations in most geological settings comparable to those found in the Netherlands did not generate M > 2.0 events. However, seismicity cannot be excluded. In particular, in faulted carbonates M > 2.0 events may occur, but the potential for these events remains site-specific and dependent on the local fault and fracture fabric and the state of stress. Taking into account the different key parameters shown in Table 1 helps to assess the possibility of the occurrence of felt events and motivate additional characterisation of the subsurface in case the potential for M > 2.0 events is larger than desired. Additionally, it is also important to inform and involve the public on the development of geothermal systems and potential hazards as we proceed with the energy transition (McComas et al., Reference McComas, Lu, Keranen, Furtney and Song2016).