geothermal engineering integrating mitigation of induced ... · this judgement may have substantial...

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Grant Agreement no. 241321-2 Geothermal Engineering Integrating Mitigation of Induced Seismicity in Reservoirs Project Acronym: GEISER

D5.1 Report on the assessment of seismic hazard associated to natural seismicity

Due date of deliverable: 1.10.2011 Actual submission date: 1.2.2012 Start date of project: 1.1.2010 Duration: 42

Lead: Swiss Seismological Service, Eidgenössische Technische Hochschule (SED-ETHZ) Partners: GFZ, BRGM, TNO, INGV Responsible: Stefan Wiemer, ETHZ Revision: 1

DisseminationLevel PU Public x PP Restricted to other programme participants (including the Commission Services) RE Restricted to a group specified by the consortium (including the Commission Services) CO Confidential, only for members of the consortium (including the Commission Services)

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1. Objectives of Task 5.1

The objectives of GEISER Task 5.1 are:

• The probability of occurrence of natural earthquakes (activity rate) is performed using

literature and catalogue data on past earthquakes and knowledge of seismo-tectonic

indicators (proximity to active faults, density of faults, stress field etc.); additional

information may be obtained from previous local geological investigations or through

pre-studies for the site assessment, including drillings, background monitoring,

modelling etc.

• The assessment of the expected ground shaking (seismic hazard) is derived combining

the earthquake occurrence probabilities with ground-motion attenuation models

calibrated for regional and local conditions

• Uncertainties in all the input parameters will be considered, to derive a complete

characterization of the resulting hazard

• A combined integrated hazard methodology for local and regional hazard assessment

will be tested on existing EGS sites

This report is revision 1 of the deliverable. It will be updated once other GEISER tasks with

links with 5.1 have been completed.

2. Introduction

Assessing the natural hazard surrounding a site is critical step in the licensing application of

EGS sites, because:

1) The added induced hazard is best communicated and judged by the public, media and

decision makers in comparisons to the natural, constant background earthquake hazard at

the site. Communication of EGS risks and perceived benefits is a core task of GEISER

WP6 and not covered here.

2) A good definition of the background hazard and the probability of natural earthquakes is

needed to assess the probability that an earthquake that occurred in the vicinity of an EGS

site is triggered by the activity. This decision may have substantial legal and financial

consequences (in terms of insurance); it also requires local and regional monitoring

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capable of resolving hypocenters with sufficient accuracy (link with Task 6.3). Below we

propose a suitable background model for EGS related activities throughout Europe.

3) Assessing when a site has returned to the natural background activity requires a good

definition of this background, ideally established before the project start. Ownership site

may for example be returned to the state or local authorities once the seismicity has

returned to the previous natural background conditions. Below we give a examples of this

assessment, leading to a suggestion on a procedure for establishing the return to

background.

4) Knowledge on the background hazard may be linked with the subsequent potential of a

site to induce and trigger earthquakes. Below we report on an the initial database efforts

and attempt to establish links between background activity and potential to induce

Below we present the main findings related to point 2, 3 and 4 in the context of the GEISER

contribution. Details of the work performed are then given in the relevant appendices to this

report.

3. Background hazard as a key to liability assessment: The SHARE community model applied for EGS background hazard assessment A good definition of the background hazard and the probability of natural earthquakes is

needed to assess the probability that an earthquake that occurred in the vicinity of an EGS site

is causally related to the activity. This judgement may have substantial legal and financial

consequences (in terms of insurance and liability); it also requires local and regional

monitoring capable of resolving hypocenters with sufficient accuracy ( WP 6). Hazard

assessment is a core responsibility of national authorities, and all nations in Europe maintain

and update a probabilistic seismic hazard model. While the degree of complexity and

sophistication varies, the basic methodology of probabilistic seismic hazard assessment

(PSHA) is well established and accepted. The GEISER WP5 team evaluated the available

models for their suitability for EGS relevant estimation of the background hazard at a

European level. We concluded that the SHARE seismic hazard model, to be published in the

summer of 2012 as part of the EC FP7 project on Seismic Hazard Harmonisation

(www.share-eu.org/) is the ideal reference framework. See box 1 for a very brief introduction

to SHARE.

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Fig 1: View of the SHARE prototype portal

The SHARE model offers the following features that make it a suitable background model for

geothermal applications:

• It is state of the art, completed in the past three years and a consensus model of

leading experts on PSHA in Europe. SHARE is a harmonised model, such that results

can be compared, and includes the most extensive knowledge on faults in Europe, a

total of 64000 km of fault sources with activity rates (Fig. 2).

• SHARE captures the epistemic and aleatory uncertainties in a comprehensive fashion.

SHARE is a living model that can be updated as new data become available; it is also

validated though time using procedures established through the European testing

model of the Collaboratory for the Study of Earthquake Predictability (CSEP).

• The model and data are open access and open source, methodology and computation

codes are well documented and fully transparent.

• Access to the SHARE community model is available through a highly interactive

portal (Fig. 1) and as WFS (web feature service). Access will be possible in the

foreseeable future as part of the EFEEHR (European Facility for Earthquake Hazard

and Risk), located at ETH Zurich and funded through the NERA (www.nera-eu.org)

FP7 infrastructure program throughout 2014.

• The SHARE community model can be extended to small magnitudes (M=1.0) needed

for geothermal applications. It is also compatible with the ground motion prediction

equations (GMPE) developed in GEISER Task 5.4.

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The SHARE and NERA teams have agreed to make a GEISER tailored instance of the

SHARE model publically available as part of EFEEHR in early 2013, in time for the release

of the GEISER ‘best practise’ tasks (5.5 and 6.5). This geothermal extension of SHARE,

maintained by EFEHR, will be a key resource for future geothermal projects in Europe.

Fig 2: Implementation of the SHARE active faults database as of 31 May 2011. The database features about 98 fully parameterized records and ~64,000 km of faults. The map shows all Composite Seismogenic Sources, color coded according to faulting mechanism: Red = normal; blue = reverse; green = strike slip and oblique slip

Box 1: A short introduction to the SHARE community model

Probabilistic seismic hazard assessment (PSHA) is one of the most useful products

seismology offers to society. PSHA characterizes the best available knowledge on the seismic

hazard of a study area, ideally taking into account all sources of uncertainty. Results form the

baseline for informed decision-making, such as building codes or insurance rates and provide

essential input to each risk assessment application.

Several large scale projects have been launched aiming to harmonize PSHA standards around

the globe. The EC-FP7 project SHARE (www.share-eu.org) released a community-based

probabilistic time-independent hazard model for the Euro-Mediterranean region in 2012 and

contributes its results to the Global Earthquake Model (GEM, www.globalquakemodel.org), a

Primary faulting style

• Normal • Reverse • Strike-slip /oblique

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public/private partnership initiated and approved by the Global Science Forum of the OECD-

GSF.

SHARE inherited knowledge from national, regional and site-specific PSHAs, assessed new

data and built comprehensive hazard relevant databases, rigorously selected the best suited

ground motion prediction equations and implemented the model in a suitable computational

framework. In addition, various contemporary applied approaches to assess sources of

uncertainties where introduced in the model within its logic-tree; for the first time kernel-

smoothed seismicity approaches as well as fault-based hazard estimates are included.

The SHARE-PSHA comprises results for various return periods of engineering interest and

various ground motion intensity measures (peak ground acceleration, spectral accelerations at

various periods) on the Euro-Mediterranean scale. In addition, details on single sites of

interest such uniform hazard spectra and disaggregation are available and accessible via an

online portal as the front end of the computational infrastructure for an integrated European

PSHA model.

SHARE is a procedural example on how to perform a regional scale PSHA addressing divers

demands from the general public, seismologists, engineers and decision makers. We present

results of the PSHA, elaborate on their implications and describe future challenges. The

results outline a foundation for future PSHAs that should include time-dependency and

sophisticated model evaluation efforts. The results can serve as the base for a European

Probabilistic Risk Assessment. We envision the results to deliver long-lasting structural

impact in areas of societal and economic relevance, to serve as reference for the revision of

Eurocode 8 (EC8) provisions and to provide a homogeneous baseline input for the correct

seismic safety assessment for critical industry, such as the energy infrastructures and the re-

insurance sector.

4. Duration of an EGS induced sequence

Once the injection of water under high pressures stops, the observed induced seismicity

decays gradually in the following days, weeks and months (Figure 3), quite similar to any

tectonic aftershock sequence. For aftershock sequences, the rate is usually well described by

the Omori-Utsu law (Ogata, 1999; Utsu, 1961):

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where t is the time elapsed after the main shock, c and p are empirical parameters,

characteristic for a specific sequence, and k(Mc) is a function of the number of events with

magnitudes above the completeness magnitude Mc.

We investigate in the context of the GEISER Task 5.1 work first of all whether the Omori-

Utsu law can provide an acceptable fit to the data of the post-injection period in Basel. We fit

the events with Mw > 0.9, that occurred between the end of the injection (December 8, at

11:33 a.m) and day 200 of the sequence. We find p = 1.33 +/- 0.06, c = 0.38 +/- 0.061 days

and k = 86.6 +/- 9.81 (Figure 3) as the mean parameters for 1000 bootstrap models. Dor

details please refer to Bachmann et al. (2011) and Bachmann (2011). The two-sample

Kolmogorov-Smirnov (Woessner et al., 2004) test, testing whether the cumulative rate of the

data and the fitted Omori-Utsu law belong to the same distribution, is not rejected at the

significance level of 0.05. This indicates a good fit of the Omori-Utsu law to our data,

suggesting that they are indistinguishable from tectonic aftershock sequences.

Figure 3: Decay of the sequence after the termination of the water injection. A modified Omori-Utsu law is fitted to the sequence to determine its duration; circles represent the data. We only fit the first 200 days of data (dark circles); events after this time (grey circles) fall within the uncertainty. The background of this region and the uncertainty are indicated at the bottom of the figure. Where they intercept with the model, we find the duration of the sequence. The black box indicates the uncertainty and the black star the best fit. We find a duration of 31 +29/−14 years (from Bachmann et al., 2011)

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To estimate the expected duration of the sequence, we additionally determine the background

seismicity rate for this region. No event has ever been located within the tiny stimulated

volume of about 1 km3 during the 25-year history of recording micro-seismicity in

Switzerland. We therefore use as a proxy the seismic activity rate of the seismogenic source

zone of Basel used in the determination of the Swiss Seismic hazard in 2004 (Giardini et al.,

2004; Wiemer et al., 2009). The rate is normalised to the size of the stimulated volume (more

specifically, to its 2D areal extension). This assumption is consistent with the definition of a

seismic source as a zone of equal seismic potential in seismic hazard assessment. The

seismogenic source zone has an a-value of 2.31, a b-value of 0.9, and spans over an area of

1741 km2. The area affected by the injection is 1.6 km2 which leads to a background rate of

Rb = 3.38e−4 events per day with Mw >= 0.9. In other words, an Mw >= 0.9 event should

occur naturally only about every eight years. If we solve the Omori-Utsu law for the duration,

we obtain (Woessner, 2005).

With these values, we obtain a duration of the ‘aftershock’ sequence of ta = 31 (+29/−14)

years, where the uncertainties are obtained from the bootstraps in Figure 3.

The same approach we then applied in Bachman (2011) to the 1993 Soultz data (Fig. 4),

finding durations of 1.15 (+1.16/-0.54) years. In the Soultz case, the completeness is Mw = -

1.0, and the estimated background 1.7 e-2 events/year.

Our assessment that the seismicity will take more than 1 year in the case of Soultz, and more

than 10 years in the case of Basel, to decay to the background after an EGS stimulation is

consistent with the observations of aftershock sequences [Stein and Liu, 2009] as well as with

predictions of laboratory studies [Dieterich, 1994]. In aftershock sequences, the duration of

the sequences, defined as the time when the rate returns to that before the event, has been

proposed to be inversely correlated with the tectonic loading rate. The rate and state model of

fault friction, which predicts changes in fault properties after earthquakes, and which is

commonly used for aftershock studies [Dieterich, 1994], predicts an aftershock duration of:

where tau is the rate of shear stressing on the fault, sigma is the normal stress, and A is a constitutive parameter. Although the stressing rate is is hard to measure, it is roughly proportional to the loading rate. The loading rate in Switzerland, and also in the Basel and Soultz regions, is known to be low (< 1 mm/yr) and aftershock sequences

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are therefore expected to last longer than in tectonically more active regions.

Figure Fig. 4: Decay after the two major injections; a) from 15.4 to 29.2 day and b) after 43.9 days. The decay between the first injection and the double packer injection (a) can not be fit with the modified Omori-Utsu law, however the decay after no additional injections follow (b) can be well fit with this law.

The major unknown in the estimation of the aftershock duration is the local background rate.

Ideally, monitoring at a comparable completeness level of the same rock volume for several

years would establish the background seismicity rate. This approach is prohibitively

expensive. We chose the regional background rate, extrapolated from the micro-seismicity

record, as a proxy, but the local variability of the background at the scale of one kilometre is

unknown. Supporting evidence for a low background rate comes from the observation that

micro-earthquakes in the past three years in Basel have been confined to the induced volume;

no events have been detected outside the well-defined volume, although the network of

borehole sensors in principle would be able to detect them.

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5. Background seismicity as an indicator for induced seismicity

Knowledge on the background hazard may be linked with the subsequent potential of a site to

induce and/or trigger earthquakes. Knowledge on such a key parameter that forecasts the

triggerability of a volume would have substantial economical consequences, as it would allow

for an a-priory assessment of the seismic response expected. Within GEISER and linked to

the GEOTHERM and CARMA projects at ETHZ, we therefore conducted an extensive site

survey, based on a dedicated database efforts conduced jointly by ETH Zurich and TNO.

Detailed results are reported in the appendix A and B to deliverable 5.1, where we also

describe in detail the database. Some of the results are also reported in Evans et al. (2011)

(Figure 5). While the existing database is too small to draw definitive conclusions, there is a

suggestion that larger magnitudes are not induced in areas of lowest hazard. As new data are

added through the remaining GEISER project, we will update the results of D5.1 accordingly.

Fig 5: Map of Europe, color-coded is the predicted peak ground acceleration (PGA) with a 10% probability of exceedance in 50 years, based on the GSHAP model. Sites studied in Evans et al. (2011) are marked; the inset in the bottom right plots the maximum observed magnitude as a function of PGA value (from Evans et al., 2011).

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WP5.1 related Presentations

Bachmann, C., S. Wiemer, B. Goertz-Allmann, J. Woessner & B. Mena: Why geothermal energy research needs statistical seismology. Statsei 7, Santorini, April 2011.

Deichmann, N.: Lessons learned from cases of induced seismicity in connection with the exploitation of deep geothermal energy. Invited talk, 6th IEAGHG Risk Assessment Network Workshop, 21–23 June, 2011, Pau, France.

Deichmann, N. and S. Wiemer: EGS and induced seismicity – the good and the bad. SGM, Zurich, Switzerland, November 2011.

Deichmann, N.: Erdbebenrisiko bei Vorhaben der tiefen Geothermie Schweizerischer Erdbebendienst. Mediation Tiefen-Geothermie, Speyer, 12/12/2011

Kraft, T., S. Husen and S. Wiemer: GEOBEST - A contribution to the long term development of deep geothermal energy in Switzerland, SGM, Zurich, Switzerland, November 2011

Goertz-Allmann, B.P., A.V. Goertz and S. Wiemer, Stress Drop of Induced Earthquakes: A Proxy for Pore Pressure? Third Passive Seismic Workshop, Athens, Greece, March 2011.

Goertz-Allmann, B.P., C. Bachmann, S. Wiemer, B. Mena, J. Woessner, and N. Deichmann: Source property variations of induced seismicity in geothermal reservoirs, Fragile Earth Conference, Munich, Germany, September 2011.

Goertz-Allmann, B.P., C. Bachmann, and S. Wiemer: What can induced earthquake source properties tell us about reservoir geomechanics? Geothermal Week, ETH-Zurich, November 2011.

Goertz-Allmann, B.P., C. Bachmann, and S. Wiemer: What can induced earthquake source properties tell us about reservoir geomechanics? SGM, Zurich, Switzerland, November 2011

Mena, B. and S. Wiemer, Building robust models to forecast the induced seismicity related to geothermal reservoir enhancement, Geothermal Week, ETH-Zurich, November 2011.

Wiemer, S., Living with induced seismicity: Lessons from Basel and a roadmap ahead, 7. Internationale Geothermiekonferenz Freiburg, 12. Mai 2011.

Wiemer S., Kleine Beben mit großer Wirkung: Ein Beitrag zur probabilistischen Gefährdungsanalyse von induzierten Erdbeben, DGG Kolloqium "Induzierte Seismizität”, Köln, 23. Feb. 2011.

Wiemer, S.: Welche Energie steckt in Erdbeben? Scientifica, ETH Zürich, August 2011.

Wiemer, S.; Corinne E. Bachmann; Bettina Allmann; Domenico Giardini; Jochen Woessner; Flaminia Catalli; Banu Mena Carbrera. A probabilistic framework for hazard assessment and mitigation of induced seismicity related to deep geothermal systems. AGU meeting, San Francisco, dec. 8 2011.

Publications related to WP5 Bachmann, C.E., Wiemer, S., Woessner, J. and Hainzl, S.: Statistical analysis of the induced Basel 2006 earthquake

sequence: introducing a probability-based monitoring approach for Enhanced Geothermal Systems. Geophysical Journal International, 186: 793–807, 2011.

Bachmann, C., S. Wiemer, B. Goertz-Allmann, and J. Woessner (2012), Influence of pore pressure on the size distribution of induced earthquakes, submitted to Geophys. Res. Lett.

Bachmann, C.E. 2011, New approaches towards understanding and forecasting induced seismicity, Ph.D Thesis, ETH Zurich.

Bethmann, F.: Magnitude scaling relations and attenuation in thick sediments: application to the induced seismicity beneath the city of Basel, Switzerland. PhD Thesis Nr. 19510, ETH Zürich, 2011.

Bethmann, F., N. Deichmann & P. M. Mai: Scaling Relations of Local Magnitude versus Moment Magnitude for Sequences of Similar Earthquakes in Switzerland. Bulletin of the Seismological Society of America, Vol. 101, No. 2, pp. 515–534, April 2011, doi: 10.1785/0120100179

Goertz-Allmann, B. P., A. Goertz and S. Wiemer: Stress drop variations of induced earthquakes at the Basel geothermal site. Geophysical Research Letters, Vol. 38, L09308, doi:10.1029/2011GL047498, 2011.

Goertz-Allmann, B.P., B. Edwards, F. Bethmann, N. Deichmann, J. Clinton, D. Fäh, and D. Giardini: A New Empirical Magnitude Scaling Relation for Switzerland, Bulletin of the Seismological Society of America, Vol. 101, No. 6, pp. 3088–3095, December 2011, doi: 10.1785/0120100291

Mena, B., S. Wiemer and C.E. Bachmann, Building robust models to forecast the induced seismicity related to geothermal reservoir enhancement, to be submitted to BSSA as a short note.

Wiemer S., Kleine Beben mit großer Wirkung: Ein Beitrag zur probabilistischen Gefährdungsanalyse von induzierten Erdbebe. DGG Kolloqium "Induzierte Seismizität", Sonderbände I/2011 der Mitteilungen der Deutschen Geophysikalischen Gesellschaft, ISSN-Nr. 091944, 14 pages

D5.1 Chapter-TNO: Key Parameter Study

Doc.nr: Version: Classification: Page:

GEISER-D5.1 Draft 2011.12.15 Public 1 of 68

All rights reserved. Copying (parts) of this document is prohibited without prior permission of TNO

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D5.1: Appendix A to deliverable 5.1

Prepared by: Janneke van der Burgt (TNO), Karin van Thienen-Visser (TNO)

Manuel Nepveu (TNO), Alba Zappone (ETH)

Reviewed by: Karin van Thienen-Visser (TNO), Manuel Nepveu (TNO)

Approved by:

Jan Hopman (TNO)





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Table of Content 1 Applicable/Reference documents and Abbreviations ............................... 3

1.1 Applicable Documents ........................................................................................................ 3 1.2 Reference Documents ........................................................................................................ 3 1.3 Abbreviations ...................................................................................................................... 3

2 Identification of key parameters related to injection-induced seismicity – Case histories for Enhanced Geothermal Systems (and other injection projects) ............................................................................................................... 4 3 Merging the ETH and TNO databases ....................................................... 66

3.1.1 Parameters to include in the merged database ......................................................... 66 3.1.2 Application to seismic hazard analysis before start of production ............................. 66 3.1.3 Database access via internet ..................................................................................... 66 3.1.4 Future Work ............................................................................................................... 67





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1 Applicable/Reference documents and Abbreviations

1.1 Applicable Documents (Applicable Documents, including their version, are the “legal” basis to the work performed) Title Doc nr Version AD-01 GEISER,

EC FP7 Grant Agreement no.: 241321 2010.01.01

AD-02 GEISER, EC FP7 Grant Agreement no.: 241321, Annex I - “Description of Work”

2009.09.28

AD-03 GEISER, Consortium Agreement, final version

2010.01.01

1.2 Reference Documents (Reference Documents are referred to in the document) Title Doc nr Version RD-01

1.3 Abbreviations (this refers to abbreviations used in this document)





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2 Identification of key parameters related to injection-induced seismicity – Case histories for Enhanced Geothermal Systems (and other injection projects)

Janneke van der Burgt(1), Dr. Karin van Thienen-Visser(2), M. Nepveu(2)

(1) Utrecht University, Faculty of Geosciences

(2) TNO: geological survey of the Netherlands, Utrecht, the Netherlands





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Abstract The major issue concerning Enhanced Geothermal Systems (EGS) is that seismic events are induced during the process of hydrofracturing aimed at increasing the permeability of the geothermal reservoir rock. Although these are predominantly microearthquakes that are not felt by humans, past injection projects have shown that magnitudes of injection-induced events can go up to 4.6.

In this study, an overview of injection histories is given for 19 EGS and other injection sites throughout Europe and from regions outside Europe. Parameters- related to the specific site, operational procedures and induced seismicity – were collected from the literature and organised in a database. These parameters were investigated with the objective to find key parameters which can provide a basis for the assessment of the seismic hazard at an EGS site prior to production.

Bayesian least squares fitting showed that the correlation between the chosen global dimensionless parameters and the maximum magnitude at a site were not good. This indicates that the underlying assumptions are not correct. Conclusions about key parameters related to injection-induced seismicity could therefore not be found. The available data does suggest that injection into granitic rocks is more prone to inducing large (felt) events in comparison to other reservoir rock. Also, the larger induced events are triggered at larger faults present in the stimulated area. Both might mean that site-specific factors are controlling the induction of large magnitude events, instead of the global dimensionless parameters used in this study.

Future plans involve both the merging of this database with the database from ETH Zurich and the extension of the database with non-geothermal sites. Data from this extended merged database will hopefully give more insight into key parameters related to injection-induced seismicity. At present, the seismic hazard at potential injection sites can only be indicated by numerical modeling of injection sites.





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4.1. Introduction Since the 1970s, the development of deep geothermal reservoirs - by the injection of large amounts of fluids at high pressures into the subsurface - has become a standard procedure. Geothermal reservoirs which need the injection of fluids to increase the permeability of the rocks, by means of hydraulic fracturing, are called Enhanced Geothermal Systems (EGS).

A major issue concerning EGS is that hydrofracturing of the subsurface rocks is often accompanied by a large amount of induced seismic events. Although most of these events have small magnitudes (they are also called microearthquakes) which are not felt by humans, some stimulations in the past have shown that relatively large events can occur (Table 1.1) [Baisch et al., 2010]. A good example is the injection at Basel, Switzerland, in 2006, which was stopped due to the injection-induced events with magnitudes up to 2.6 [Häring et al., 2006]. Several hours after shut-in of the well, an even larger event took place having a magnitude 3.4 [Häring et al., 2006]. The stimulation of the 5 km deep geothermal reservoir at Soultz-sous-Forêts in France also induced relatively large events with magnitudes up to 2.9 [Dorbath et al., 2009]. Since the number of geothermal projects is increasing in industrialised and residential regions, it is very important to assess the seismic hazard of a potential EGS site before the start of production.

Van Eijs et al. (2006) investigated the seismic hazard for gas depletion before the onset of seismicity by using data from hydrocarbon sites in the Netherlands. In this study they looked at the correlation between several parameters related to reservoir and production properties and the occurrence of induced seismicity during gas depletion. They found three key parameters which showed a good correlation with the occurrence of induced events: pressure drop at first earthquake, the stiffness contrast between the reservoir rock and the overburden and reservoir fault density.

Although many studies have been done on induced seismicity related to EGS, there is still little information available on how induced seismicity can be predicted before the onset of production at a geothermal site. This paper gives an overview of the injection history of 19 sites throughout Europe and from regions outside Europe in order to obtain physical parameters with predictive properties for induced seismicity associated with EGS operations. This study is part of a large European project called GEISER (Geothermal Engineering Integrating Mitigation of Induced Seismicity in Reservoirs) and may lead to a better understanding of these key parameters, providing a basis for assessing the seismic hazard at future EGS sites.

First, the project histories of the studied sites are described. Then only a few potential parameters, selected from the database, are plotted as a function of the maximum magnitude at each site with the aim to find parameters which show a correlation with the maximum magnitude induced at injection sites.

Recently, Evans et al. (in press) studied the injection-related seismicity of sedimentary and crystalline rocks at 41 European sites in order to identify factors which help to assess the seismic hazard related to fluid injection. Contrary to prior concepts, they found that injections at larger depths in the crystalline rocks did not induce larger events than those at shallower depth. Also, they suggested that low values for the local peak ground acceleration (natural seismicity) (< 0.07 g) could be an indicator for the induction of small events during fluid injection. However, they found that this does not mean that high values for the peak ground acceleration result in the induction of a higher level of seismicity.

Site Mmax Site Mmax Soultz, GPK-1 (1993) 2.0 The Geysers (1965-present) 4.6 Soultz, GPK-2 (2000) 2.5 Cooper Basin (2003) 3.7* Soultz, GPK-3 (2003) 2.9* Cooper Basin (2005) 2.9 Soultz, GPK-4 (2004) 2.3* Berlín (2002-2003) 4.4* Soultz, GPK-4 (2005) 2.7 Ogachi (1991) 2.0

Basel (2006) 3.4* Paradox Valley (Phase I) (1996-1999)

3.6





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KTB (1994) 1.2 Paradox Valley (Phase II) (1999-2000)

4.3

KTB (2000) 0.7 Paradox Valley (Phase III) (2000-2002)

2.8

Groβ Schönebeck (2007) -1.1 (Mw) Latera-1 (1981 & 1982) 0.5

Latera-6 (1981) 0.4 Table 1.1 Overview of maximum local magnitudes (Mmax) for different injection projects. *occurred after shut-in 4.2. Overview of project histories In this section, the injection histories of the studied sites are described, starting with those included in the GEISER project and then the other injection projects encountered in the literature are reviewed. The sites included in GEISER are mostly located in Europe (Figure 2.1). Detailed information on all the parameters, for the sites described below, is listed in the database in Appendix I. These parameters include those related to the geology (e.g. fracture network, depth of injection/reservoir, type of reservoir rock, temperature, faults), seismicity (e.g. maximum magnitude, amount of events), stress (e.g. magnitude of minimum and maximum horizontal stress), pressure (e.g. maximum formation pressure) and injection (e.g. volume injection, maximum injection rate, maximum wellhead pressure) at each site. Definitions of the parameters listed in this database are given in Appendix II. 4.2.1 Sites included in the GEISER project 4.2.1.1 Soultz-sous-Forêts (France)

The Soultz-sous-Forêts geothermal site is located in the Upper Rhine Graben (France) and has a granitic reservoir with temperatures reaching ~160° C at 3.5 km and ~200° C at 5 km depth (Figure 2.1) [Majer et al., 2007; Dorbath et al., 2009]. The granite is pervasively fractured with natural fractures in the granite generally striking N160° E, subparallel to the maximum horizontal stress [Baria et al., 1999; Valley, 2007; Dezayes et al., 2010]. All of the seismic clouds which formed during the stimulations at Soultz are subvertical planar structures oriented N20°W to N-S and are thus subparallel to the maximum horizontal stress [Evans et al., 2005; Dorbath et al., 2009].

The first well (GKP-1) was drilled to about 2,000 m depth in 1987 and was subsequently deepened to about 3,600 m in 1992 [Tester et al., 2006]. In September 1993 ~25,000 m3 of water was injected in the open-hole section at a maximum injection rate of





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Figure 2.1 Map showing the locations of the studied EGS sites in Europe (Google Earth).

36 l s-1 [Evans et al., 2005]. About 12,000 induced events were recorded during the injection of which the largest event had a magnitude of 2.0 [Evans et al., 2005; Deichmann and Evans, 2010]. The hypocenter locations form a diffuse, subvertical cloud striking N25°W without apparent internal structure [Evans et al., 2005].

In 1995, a second borehole (GPK-2) was drilled to 3,876 m depth [Tester et al., 2006]. GPK-2 was hydraulic stimulated in 1995 and 1996 by injecting 28,000 m3 (in each injections) with maximum injection rates of 56 and 78 l s-1, respectively [Baria et al., 1999; Baisch and Vörös, 2009]. After GPK-2 was deepened to 4,955 m, the well was stimulated in 2000 by injecting 22,680 m3 of water for six days with a maximum injection rate of 50 l s-1 [Dorbath et al., 2009]. With the onset of injection, event rates were immediately very high and declined very rapidly after shut-in [Oates et al., 2004]. More than 30,000 events were recorded with the largest event (during the injection) having a magnitude of 2.5 [Dorbath et al., 2009]. Six days after the start of injection and 9 hours after shut-in, the second largest event with magnitude 2.4 occurred [Charlety et al., 2007]. The largest event (M=2.6) of this stimulation took place 10 days after shut-in [Charlety et al., 2007]. The hypocenter locations of the induced events were clustered around the injection well





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Figure 2.2 The hypocenter locations of the induced events from the stimulation at Soultz in 2000: a) Vertical cross-section through the seismic cloud (containing events with M > 1) along a line N20°W. b) Vertical cross section through the seismic cloud (containing events with M>1) along a line N70°E [source: Dorbath et al., 2009].





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and migrated away from the injection well during the stimulation [Oates et al., 2004]. The hypocenter locations formed a dense, homogeneous cloud without evident internal structures and striking N20°W (Figure 2.2) [Dorbath et al., 2009].

In 2002-2003, GPK-3 was drilled to a depth of 5,091 m and it was stimulated in 2003 by injecting 37,300 m3 for about eleven days with a maximum injection rate of 90 l s-1 [Majer et al., 2007; Dorbath et al., 2009]. 90,648 events were recorded and the largest event (M=2.9) occurred after stimulation had ended, also being the largest event for all the stimulations at Soultz [Dorbath et al., 2009]. In contrast to the 2000 stimulation, the seismic cloud was sparse and showed some large north-south structural features dipping steeply to the west [Dorbath et al., 2009]. This suggests that most of the events occurred on large faults, which would explain the larger events (more area available to slip) [Dorbath et al., 2009]. The triggered events on nearby faults explain why the amount of medium size earthquakes (M<2) was lower and the number of large earthquakes (M>2) was higher in comparison to the 2000 stimulation of GPK-2 [Dorbath et al., 2009]. Also, the b-value of 0.94 from the Gutenberg-Richter plot for the stimulation of GPK-3 is comparable to b-values in areas dominated by fault tectonics [Dorbath et al., 2009]. Seismic activity during this injection was independent of the injection rate and solely determined by the dynamics of the fault [Dorbath et al., 2009].

GPK-4 was drilled down to 4,982 m in 2004 and has been stimulated twice [Dorbath et al., 2009]. In September 2004, 9,300 m3 of water was injected with a maximum velocity of 30 l s-1 and in February 2005 12,300 m3 water with a maximum velocity of 45 l s-1 was injected [Dorbath et al., 2009]. During both stimulations 22,718 induced events were recorded with the largest event having a magnitude of 2.3 and 2.7 for the 2004 and 2005 stimulation, respectively [Dorbath et al., 2009]. Just as in the 2003 stimulation, the magnitude 2.3 event occurred after stimulation [Baisch et al., 2010]. The distribution of seismicity suggests that most of the events took place on a single fault zone striking approximately N-S and steeply dipping to the west [Dorbath et al., 2009].

4.2.1.2 Basel (Switzerland)

The Basel geothermal site is situated in the southern Upper Rhine Graben (Switzerland) and consists of 1 borehole (Basel-1) (Figure 2.1). Basel-1 was drilled to about 5 km depth, encountering 2.4 km of sedimentary rocks lying on top of 2.6 km of granite basement [Häring et al., 2008]. Temperatures are similar to those found at Soultz with 200° C at 5 km depth [Häring et al., 2007]. Initial permeability was very low for the Basel reservoir, suggesting that the granitic reservoir at Basel is not pre-fractured like the Soultz reservoir [Häring et al., 2007].

The background seismicity in the Upper Rhine Graben is characterised by frequent minor seismicity [Häring et al., 2008]. However, occasionally destructive seismicity does occur in this area; in 1356 the largest known historical earthquake in central-northern Europe, with an estimated local magnitude between 6.5 and 6.9, occurred in Basel [Häring et al., 2008]. Most of the naturally occurring earthquakes have been small to moderate events in recent times [Häring et al., 2008].

Hydraulic stimulation of Basel-1 was performed in December 2006 by the injection of 11,500 m3 water. During six days of injection maximum injection rates were 60 l s-1 [Asanuma et al., 2007; Häring et al., 2008; Dinske, 2010]. Dinske (2010) found that the b-value was decreasing with time of injection, and explains this by the expansion of the stimulated reservoir leading to reactivation of larger fracture planes [Dinske, 2010]. Stimulation was stopped on December 8th due to the occurrence of an event having a magnitude 2.6 [Baer et al., 2007]. However, four hours after shut-in, an even larger event of magnitude 3.4 occurred [Häring et al., 2007]. Kahn (2008) suggests that this magnitude 3.4 event took place on a vertical plane oriented N65°W.

During injection more than 10,500 events were recorded [Baer et al., 2007]. The locations of the induced events formed a subvertical elongated zone of seismic activity striking subparallel to the direction of SHmax (N160° E) and thus also subparallel to the majority of natural fracture orientations (140-170°) (Figure 2.3) [Asanuma et al., 2007; Dyer et al., 2008; Dinske, 2010]. This shows that the preferred direction of fluid migration is in the direction of SHmax [Dinske, 2010]. The





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shape and orientation of this seismic cloud is very similar to the form of the clouds during the stimulations at Soultz-sous-Forêts, France.

Figure 2.3 Hypocenter locations of the main stimulation at Basel (2006) (source: Häring et al., 2007].





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4.2.1.3 KTB borehole (Windischeschenbach, Germany)

Two deep geothermal wells were drilled as part of the German Deep Drilling Program KTB (Kontinentales Tiefbohrprogramm der Bundesrepublik Deutschland) with the goal to study the properties and processes of the deeper continental crust [Emmermann and Lauterjng, 1997]. The KTB site is located at the western margin of the Bohemian Massif in NE Bavaria (Germany) where natural seismicity is low (Figure 2.1) [Baisch et al., 2002; Zang and Stephansson, 2010]. The wells were drilled in a small tectonometamorphic unit called the Zone of Erbendorf-Vohenstrauβ (ZEV) at the boundary between the Moldanubian and the Saxothuringian units [Emmermann and Lauterjung, 1997]. The Franconian Lineament, a NW-SE trending system of reverse faults, separates the basement rocks from the Permo-Mesozoic foreland sediments (Figure 2.4) [Zang and Stephansson, 2010]. Orientation of the maximum horizontal stress at KTB is about 160º, which is similar to the average orientation found throughout western Europe [Brudy et al., 1997].

Figure 2.4 Vertical cross-section through the KTB-HB well [source: Zang and Stephansson, 2010].

Between 1987 and 1990 the pilot hole, KTB-VB (VorBohrung), was drilled reaching a

depth of 4,000 m [Zang and Stephansson, 2010]. The main hole, KTB-HB (HauptBohrung), was drilled to a depth of 9,101 m at 200 m distance from the pilot hole [Zang and Stephansson, 2010]. The planned depth of the well (10-12 km) was not reached due to the bottom hole temperature of 265º C [Zang and Stephansson, 2010]. The entire KTB-HB well consists of steeply inclined units belonging to ZEV, which contains mainly gneisses and metabasic rocks [Zang and Stephansson, 2010]. Major fault systems, consisting of a number of individual fault planes mainly dipping ~40º to the NE, intersect the KTB-HB well (Figure 2.4) [Bohnhoff et al., 2004; Zang and Stephansson, 2010]. These fault systems penetrate the well between 6,850 and 7,260 m depth (SE1), at 4 km depth (SE2) and between 7,820 and 7,950 m (Figure 2.4) [Zang and Stephansson, 2010].





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Injection of 210 m3 of heavy brine into the open-hole section of the KTB-HB well took place in 1994 [Zang and Stephansson, 2010]. This short-term injection (24 hours), with injection rates up to 9 l s-1 and maximum well head pressures of 50 MPa, induced 400 events [Zang and Stephansson, 2010]. Hypocenter locations are mostly clustered near the open-hole section (bottom of the well) forming an SW-NE elongated cloud [Baisch et al., 2002]. A second cluster of hypocenters was present at 8.5 km depth, which was probably due to leakage in the casing [Baisch et al., 2002]. The largest induced event had a local magnitude of 1.2 [Baisch and Vörös, 2009].

In 2002, long-term injection took place with 4,000 m3 of fresh water injected in 60 days into the open-hole section of the KTB-HB well [Baisch et al., 2002]. Injection rates were between 0.5 and 1.2 l s-1 and the maximum wellhead pressure was 30 MPa [Baisch et al., 2002]. During this long-term injection about 2,799 events were recorded, which were mainly clustered near the main open-hole between 5 and 6 km depth due to leakage of the casing [Baisch et al., 2002]. This seismic cloud was elongated in the SE-NW direction [Baisch et al., 2002]. A smaller cluster of hypocenter was located at 8.0-9.2 km depth forming a bowed structure striking N-S [Baisch et al., 2002]. When comparing these hypocenter locations with those from 1994, it is clear that only a few events were located in the same locations as in 1994 (Figure 2.5) [Baisch et al., 2002]. The events from the 2000 injection were mostly located at the rim of the seismically stimulated volume of 1994 [Baisch et al., 2002]. The largest event induced during the 2000 injection had a magnitude of 0.7 [Baisch and Vörös, 2009].

The number and maximum magnitude of induced events during these two stimulations were significantly lower than those recorded during the stimulations at Soultz-sous-Forêts and Basel.

Figure 2.5 Map view showing the hypocenter locations for both the 1994 and 2000 stimulation at KTB [source: Baisch et al., 2002] 2.1.4 Groβ Schönebeck (Germany)

The Groβ Schönebeck geothermal site is situated in the Northeast German Basin, where the background seismicity is negligible (no recent tectonic activity) (Figure 2.1) [Kwiatek et al,. 2010]. The geothermal reservoir consists of Lower Permian sandstones and volcanic rocks (Andesites) with temperatures up to 150° C and is overlain by a Zechstein salt formation [Kwiatek





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et al., 2010; Moeck et al., 2009]. The sandstones are permeable, whereas the permeability of the volcanic rocks is mainly due to the presence of natural fractures [Zimmermann et al., 2009]. There are two types of faults present in the reservoir, normal and strike-slip faults [Moeck et al., 2009]. The normal faults strike NE-SW and dip ~50° to the SE or NW and the strike-slip faults strike both NE-SW and NW-SE and have a dip larger than 80° [Moeck et al., 2009].

Two wells were drilled at this site: E Grsk 3/90 and Gt GrSk 4/05 [Kwiatek et al,. 2010]. Hydraulic stimulation took place in the GrSk 4/05 well, which is inclined at 47° in the reservoir section and extends to ~4,285 m TVD [Kwiatek et al., 2010; Zimmermann et al., 2010]. This well was drilled in the direction of the minimum horizontal stress (~108.5°) to ensure orientations of hydraulically induced fractures in the direction of the maximum horizontal stress (~018.5°) [Moeck et al., 2009; Kwiatek et al., 2010].

Starting on August 9th 2007, the volcanic section was stimulated for 6 days by injecting 13,000 m3 of water with the addition of 24 tons of sand acting as a proppant [Kwiatek et al., 2008]. During this waterfrac treatment, injection rates were up to 150 l s-1 at injection pressure of 58.6 MPa [Kwiatek et al., 2008]. Although injection rates were very high, only ~70 seismic events were detected during this stimulation with a maximum moment magnitude of -1.1 [Kwiatek et al., 2008]. About 20 events (sequence A) occurred ~20 minutes after the start of the first injection test and on August 13th another seismic sequence (sequence B) occurred after a significant decrease in flow rate (Figure 2.6) [Kwiatek et al., 2010]. After shut-in of the well on August 14th, the last seismic sequence (sequence C) occurred (Figure 2.6) [Kwiatek et al., 2010]. The hypocenter locations of the clusters B and C lie in a plane with a strike of 17° and a dip of 52° SE (Figure 2.6) [Moeck et al., 2009]. Moeck et al. (2009) suggested that these events took place on a normal fault.

On the 18th and 19th of August the sandstone formation was stimulated twice by injecting 500 m3 of crosslinked gel with 100 tons proppants in each of the treatments [Kwiatek et al., 2008]. The sections between 4,005 and 4,009 m TVD and 4,068-4,070 m TVD were stimulated with injection rates of 66 and 58 l s-1, respectively [Kwiatek et al., 2008; Moeck et al., 2009]. Maximum well head pressures were 49.5 and 38 MPa, respectively [Kwiatek et al., 2008]. In contrast to the stimulation of the volcanic rocks, almost no seismic events were induced during stimulation of the sandstones [Moeck et al., 2009].

It was not expected that seismicity would be low (both in number and maximum magnitude recorded) at the Groβ Schönebeck site due to the high injection rates and wellhead pressures [Urpi et al., 2008; Kwiatek et al., 2010]. Especially considering the earlier experience of stimulation of granitic reservoirs at Soultz-sous-Forêts and KTB which induced thousands of induced events with lower injection rates and injection pressure [Kwiatek et al., 2010]. This suggests that the local geologic and tectonic setting are of great importance for the induction of seismicity [Kwiatek et al., 2010]. The temporal behaviour of seismicity at Groβ Schönebeck was similar to the injections in crystalline reservoir with the majority of events occurring at the end of major stimulation phases, after maxima in injection rate and formation pressure [Kwiatek et al., 2008].

The overall low seismicity (rates and magnitudes) can be explained by the difference in shear strength; the shear strength of sedimentary rocks is lower than for crystalline rocks [Kwiatek et al., 2008]. Also, the presence of the Zechstein salt formation lying on top of the reservoir makes it harder to record events with recording equipment at the surface or at shallow depth, which might explain the low amount of seismicity recorded [Urpi et al., 2011]. Low stress conditions at the reservoir depth for the volcanic section can explain the small number of induced events during stimulation of the volcanic rocks [Moeck et al., 2009]. However, it is harder to explain why almost no seismicity was induced during the stimulation of the less competent and highly stressed sandstones [Moeck et al., 2009]. Moeck et al. (2009) proposed that the larger volumes injected at higher pressures into the volcanic section, in comparison to the stimulation of the sandstones, might explain this [Moeck et al., 2009].





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Figure 2.6 Perspective view of the E Grsk 3/90 and Gt GrSk 4/05 wells and the hypocenter distribution for the induced events during the stimulation of well Gt GrSk 4/05 in 2007. The plane represents the least squares fit to the locations of events from clusters B and C1. Hypocenter locations are also projected on two sides of the rectangular box. The color scale on the right represents the hypocentral depth of the plotted events, for reference the wells are given the same colors corresponding to the different depths [source: Kwiatek et al., 2010].





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4.2.1.5 Rosemanowes (UK)

At the Rosemanowes site, located in Cornwall (UK), stimulation experiments took place in the natural jointed Carnmenellis granite (Figure 2.1) [Parker, 1999]. The two main joint sets are subvertical and strike 150° N (most populated) and 70° N [Bruel, 1995]. The maximum horizontal stress strikes 130° N, differing approximately 20° from the most populated fracture orientation [Bruel, 1995].

Research at Rosemanowes took place in three phases. Phase I (1977-1980) involved the drilling of 300-m deep boreholes to test some possible fracture-initiation techniques [Richards et al., 1992; Tester et al., 2006]. Investigation of reservoir development at around 2 km depth took place in Phase 2 (1980-1988) [Parker, 1999]. In 1980 (begin of Phase 2A) a production well (RH11) and an injection well (RH12) were drilled to about 2,000 m depth with bottom hole temperatures of 79º C [Parker, 1999]. The lower sections of these wells are inclined at an angle of 30º (to the NW) from vertical [Parker, 1999]. Parker (1991) states that for a producible geothermal reservoir the wells should have been drilled in the NE direction to intersect the most populated joints. Stimulation of the natural joints was carried out in 1982 by injection of 30,000 m3 water into RH12 with maximum injection rates of 100 l s-1 and maximum wellhead pressures of 14 MPa [Parker, 1991; Parker, 1999].

Figure 2.7 Location of induced events due to the viscous stimulation of RH15 in July 1985 [source: Parker, 1999].





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Phase 2B (1983-1986) started with drilling of a production well (RH15) on a spiral track to 2,600 m TVD with bottom hole temperatures of ~100º C (Figure 2.7) [Parker, 1999; Tester et al., 2006]. The orientation of this well was much more favourable than the previous two wells, because it was drilled at right angles to the orientation of RH11 and RH12 [Parker, 1999]. This means that RH15 was also oriented at right angles to the major natural joint set through which flow was taking place [Parker, 1999]. Viscous stimulation of the upper part of the open-hole section of RH15 took place in July 1985 and lasted for 8 hours [Green and Baria, 1987]. In this stimulation 5,500 m3 of viscous gel was injected at an average flow rate of 198.2 l s-1 and a maximum wellhead pressure of 15 MPa [Green and Baria, 1987]. After the gel injection, 200 m3 of water was injected with a maximum wellhead pressure of 16.3 MPa [Green and Baria, 1987]. During the stimulation, 270 induced events were detected forming a vertical tabular structure of about 980,000 m3 (Figure 2.7) [Green and Baria, 1987]. These events were caused by shearing of natural joints in the granite [Green and Baria, 1987]. Most of the seismicity occurred during the first 2 hours, after that the event rate became lower [Green and Baria, 1987].

During the remaining period of Phase 2B and Phase 2C (1986-1988) long-term circulation took place in the reservoir created during Phase 2B [Parker, 1999]. In June 1990, RH15 (during Phase 3A) was stimulated again by injecting 4,000 m3 of low viscosity gel followed by 40 m3 of high viscosity gel and then 20 m3 of gel slurry with 11 tonnes of bauxite proppant [Parker, 1999]. During this stimulation, a relatively large volume of rock was stimulated next to the reservoir created in Phase 2B and was oriented vertically [Parker, 1999].

4.2.1.6 Reykjanes, Hengill, Krafla (Iceland)

At Iceland high temperatures are present at relatively shallow depth, because of the Mid-Atlantic ridge, resulting in active rifting and volcanism (Figure 2.8) [Friðleifsson and Elders, 2005]. Additional heat is provided by the mantle plume located beneath Iceland [Friðleifsson and Elders, 2005]. These geological circumstances make Iceland a favourable place for geothermal energy as a source of power [Kahn, 2008]. Micro-earthquakes are very common at Iceland [Kahn, 2008]. 4.2.1.6.1 Reykjanes

The Reykjanes site is one of the smallest geothermal fields in Iceland and high reservoir

temperatures, ranging between 260 and 310º C, are probably due to an active sheeted dike swarm [Sveinbjornsdottir et al., 1986; Friðleifsson and Elders, 2005; Kadko et al., 2007]. This geothermal field is located in the southwest of Iceland, in an immature rift system (Figure 2.8) [Friðleifsson and Elders, 2005]. The Reykjanes peninsula consists of young and high permeable rock formations and is tectonically active [Sveinbjornsdottir et al., 1986]. Due to this high permeability and the low level elevation of Reykjanes, sea water can percolate into the bedrock [Sveinbjornsdottir et al., 1986]. Therefore geothermal wells in this area, drilled to depths from 1,000 to over 2,500 m, produce altered sea water [Friðleifsson and Elders, 2005; Kadko et al., 2007]. The first 1,000 m of these drilled wells consists of hyaloclastic tuffs, breccias and tuffaceous sediments, whereas the lower part comprises of basaltic lavas and tuffaceous rocks [Sveinbjornsdottir et al., 1986]. 4.2.1.6.2 Hengill

The Hengill volcanic system also lies in the southwest of Iceland, about 30 km east of

Reykjavik and is one of the largest geothermal fields in Iceland (Figure 2.8) [Friðleifsson and Elders, 2005; Franzson et al., 2010; Hardarson et al., 2010]. The background seismicity at this area is very high, because the Hengill volcano is located at a triple junction where two active rift zones and a seismically active transform zone meet [Hardarson et al., 2010]. The Hengill geothermal field is situated in a tectonically active graben in which major fractures and faults





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strike NE-SW [Hardarson et al., 2010; Sanjuan et al., 2010b]. Permeability is mostly controlled by fractures [Sanjuan et al., 2010b]. Reservoir temperatures vary between 200 and 320º C and the system is recharged by meteoric water [Franzson et al., 2010; Sanjuan et al., 2010b].

Figure 2.8 Map of Iceland showing zones of active rifting and volcanism. The Reykjanes and Hengill (Nesjavellir) geothermal fields are located in the southwest, whereas the Krafla geothermal field is situated in the northeast [source: Friðleifsson and Elders, 2005].





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The Hengil geothermal area can be subdivided into four geothermal fields: Nesjavellir, Hellisheiði, Bitra and Hverhalið [Franzson et al., 2010]. A total of 90 wells, including deep exploration, production and reinjection wells, have been drilled at Nesjavellir and Hellisheiði [Franzson et al., 2010]. At the other two geothermal fields, only a few exploration wells have been drilled [Franzson et al., 2010]. In 1965, the first five exploration wells were drilled at Nesjavellir [Franzson et al., 2010]. First exploration of the geothermal resource at Hellisheiði started in 1986, but main drilling commenced in 2001 [Franzson et al., 2010]. Large fault structures located at the western edge of the Hengill graben have been targeted for drilling at the Hellisheiði field, because they act as major permeable structures [Franzson et al., 2010]. All of these wells range in depth between 1,000 and 3,322 m (deepest well drilled in Iceland) [Franzson et al., 2010]. The wells drilled at the Nesjavellir field are mostly vertical, whereas those drilled in the newer geothermal fields are directionally drilled [Franzson et al. 2010]. At Nesjavellir, hyaloclastite accumulations are dominantly present down to about 400 m, whereas lava accumulates dominate below this depth [Franzson et al., 2010]. At Hellisheiði, the transition between hyaloclastites to lava accumulates occurs at a slightly larger depth (at ~800-1,000 m ) [Franzson et al., 2010].

At Hellisheiði all of the waste water is reinjected, which was at first a method solely aimed at disposing of wastewater from power plants [Hardarson et al., 2010]. However, it was quickly discovered that this procedure also counteracted the pressure drop due to production [Hardarson et al., 2010]. The downside of injection of fluid into the reinjection wells at Hellisheiði is that it resulted in seismicity in the area around the well [Hardarson et al., 2010]. Reinjection is now obligatory by law and has caused several events since. The largest event associated with reinjection of produced water occurred 15 October 2011 in the Gráuhnúkar area. 4.2.1.6.3 Krafla

The Krafla geothermal field (northern part of Iceland) consists of several subfields and is situated in an active volcanic caldera, directly above a mantle plume (Figure 2.8) [Sveinbjornsdottir et al., 1986; Kahn, 2008]. This caldera is cut by a long fissure swarm [Ármannsson et al., 1989]. Separation associated with the Mid-Atlantic Ridge is especially noticeable at this geothermal field; between 1975 and 1984 displacements due to rifting had a total length of almost 7 meters [Kahn, 2008].

The first geothermal wells were drilled in 1974 [Ármannsson et al., 1989]. At present, there are 20 production and 2 reinjection wells in this area, ranging between 1,000 and 2,400 m and producing meteoric water with a small amount of volcanic gases [Bertani, 2005; Friðleifsson and Elders, 2005; Sanjuan et al., 2010b]. The geothermal reservoir consists of a shallow single-phase liquid reservoir (300-1,200 m) and a deep two-phase reservoir (1,000-2,000 m) having temperatures of 190-210º C and 240-340º C, respectively [Bodvarsson et al., 1984a; Bertani, 2005]. The reservoir consists of volcanic rocks with sequences of basalt flows, hyaloclastics and intrusions and fluid flow in the reservoir is fracture dominated [Bodvarsson et al., 1984a].





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4.2.1.6.4 Iceland Deep Drilling Project (IDDP)

The presence of supercritical fluids was discovered in 1985, in well NJ-11 (Nesjavellir) [Friðleifsson and Elders, 2005]. This discovery led to the IDDP (Iceland Deep Drilling Project), a long-term project which aims at producing supercritical fluids from wells drilled to depths between 3.5 and 5 km and temperatures of 450-600º C [Friðleifsson and Elders, 2005]. The Krafla, Hengill (Nesjavellir) and Reykjanes geothermal fields were selected for drilling three deep geothermal wells [Friðleifsson and Elders, 2005]. As described above, it is very common for temperatures to exceed 300º C at wells drilled to only 2 km depth in these areas [Friðleifsson and Elders, 2005]. Frequent seismic activity below 5 km at these sites indicates that the rocks are brittle at large depths and thus likely permeable, even at high temperatures [Friðleifsson and Elders, 2005]. In 2009, drilling of the first well IDDP-1, to a target depth of 4,500 m, commenced at Krafla [Friðleifsson et al., 2010]. However, drilling was terminated at a depth of 2,104 m when magma was encountered there [Friðleifsson et al., 2010]. Although the initial goal was to drill a deeper well, superheated steam can possibly be produced from the magma by injecting water into nearby wells and thereby creating the World’s hottest EGS system [Elders and Friðleifsson, 2010]. Drilling of the other two wells IDDP-2 (Hengill) and IDDP-3 (Reykjanes) to a depth of 4 km will take place in 2011-2012 and will then subsequently deepened to reach supercritical fluids [Elders and Friðleifsson, 2010]. If the outcome of this project shows that the economics of geothermal energy is improved by using supercritical fluids, this might also be applied to other geothermal fields [Friðleifsson and Elders, 2005]. 4.2.1.7 Latera (Italy)

The Latera geothermal field is located in the Vulsini volcanic complex in Latium (Italy) (Figure 2.1) [Cavarretta et al., 1985]. The geothermal reservoir consists of fractured metamorphosed limestone and igneous rocks [Cavarretta et al., 1985]. The seismic response of the reservoir was investigated during injection in the Latera-1 and Latera-6 wells. All the information on these injections, used in the description below, is taken from the data distributed among the GEISER partners.

Between June 1981 and May 1982, three injection experiments were performed in the Latera-1 well. This well has a depth of ~2,783 m and a bottom hole temperature of ~350° C [figure 6 in Cavarretta et al., 1985]. In June 1981, injection was performed at maximum injection rates of ~17 l s-1 and maximum wellhead pressures of ~7 MPa for about two days. The 2nd injection took place during 4 days in February 1982 with a similar maximum wellhead pressure and a slightly higher maximum injection rate of about 26 l s-1. Much higher maximum injection rates of 125 l s-1 were used for the third injection interval. This injection took place between 13

and 20 May 1982, but injection was not continuous; the total injection time was only about 43 hours. The level of induced seismicity was very low during these three injections. During the first injection, 193 events were recorded with the largest event having a magnitude of -0.4. During the second and third injection 147 and 372 events were recorded, respectively. Maximum magnitudes were 0.2 and 0.5 for the second and third injection, respectively. In December 1981, three injections were carried out in the Latera-6 well. Latera-6 has been drilled to a depth of ~2,017 m where temperatures are about 225° C [figure 6 in Cavarretta et al., 1985]. The duration of the first injection was about 16 hours during which fluid was continuously injected at a rate of about 67 l s-1. During the second injection, lower maximum injection rates (~28 l s-1) were used and the maximum wellhead pressure was about 15 MPa. Injection was not continuous in this injection; total injection time was about 40 hours between 16 and 19 December. Between 21 and 22 December the third injection was performed with a maximum wellhead pressure of ~14 MPa at a continuous injection rate of about 40 l s-1. Again, microseismicity was very low during the injection into Latera-6. Only 20 events were recorded during the first injection, having only negative values going up to -0.4. 90 events were recorded





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during the second injection with also only negative values and a slightly higher maximum magnitude of -0.1. The third injection induced more events, ~193 events were recorded with a maximum magnitude of 0.4. 4.2.1.8 Campi Flegrei (Italy) Campi Flegrei, situated in the Naples area (southern Italy), is an active volcanic area with as main structure a large volcanic caldera (Figure 2.1) [De Natale and Troise, 2011]. The presence of a large caldera in combination with large aquifer systems makes Campi Flegrei an ideal location for the production of geothermal energy [De Natale and Troise, 2011]. Several geothermal exploratory wells have already been drilled to depths of ~3 km at the Mofete and San Vito geothermal fields [De Vivo et al., 1989]. The San Vito field is located in the center of the caldera, whereas the Mofete field is situated in the western part [De Vivo et al., 1989]. Measured temperatures at depths between 2.5 and 3 km are between 350 and 420º C in these wells [Valentino et al., 1999].

Figure 2.9 Sketch of a vertical cross-section through the planned CFDDP well at Campi Flegrei, showing the likely substructure which will be sampled by this deep well [source: de Natale and Troise, 2011].





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Future plans at Campi Flegrei include drilling of a 4 km deep well into the caldera as part of the ‘Campi Flegrei Deep Drilling Project’ (CFDDP) (Figure 2.9) [De Natale and Troise, 2011]. The drilled well should encounter temperatures of around 500-600º C at its maximum depth [De Natale and Troise, 2011]. It is expected to find supercritical fluids at depth, which become superheated steam at the surface, being able to generate very high electrical powers [De Natale and Troise, 2011]. One of the main aims of the CFDDP is to directly analyse the geothermal system along its entire depth in order to understand the best way to exploit geothermal energy at this region [De Natale and Troise, 2011]. So far, such a drilling project has only been planned for Iceland (section 2.1.6.4) [De Natale and Troise, 2011].

Figure 2.10 Locations of EGS (green) and other injection sites (purple) in the US (Google Earth). 4.2.1.9 The Geysers (California, US) The Geysers geothermal field is situated in California, about 193 km north of San Francisco and is currently the largest geothermal field in the world (Figure 2.10) [Karner, 2005; Mossop and Segall, unpublished study]. This site has 43 reinjection wells and 424 production wells at reservoir depths between 600 and 3,000 m [Bertani, 2005]. The naturally fractured reservoir rock consists of greywacke, which has a low permeability, and is underlain by a silicic batholith (felsite) [Lipman et al., 1978; Karner, 2005]. Fractures in the greywacke are randomly oriented and subhorizontal, whereas those in the felsite are NNW oriented and near vertical [Karner, 2005]. The reservoir contains dry steam and the initial temperature before production





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was about 240° C [Lipman et al., 1978; Denlinger and Bufe, 1982]. Steam was produced from both the fractures in the greywacke and the felsite [Mossop and Segall, unpublished study]. The area surrounding The Geysers is cut by many faults and has a high seismic activity [Majer and Peterson, 2007]. Two large faults just outside The Geysers area are the Mayacamas Fault (active) and the Collayomi Fault (inactive) [Majer and Peterson, 2007]. In the last 10,000 years, no activity has taken place on the faults in The Geysers field [Majer and Peterson, 2007].

Production of steam started in the 1960s, production rates increased from 9 million kg steam per day in 1966 to about 92 million kg steam per day in 1974 [Denlinger and Bufe, 1982]. This increase in production rate was accompanied by an increase in seismicity, which Denlinger and Bufe (1982) proposed to be due to the decrease in temperature and pressure. Water was simultaneously injected to prevent the rapid decline of reservoir pressures and flow rates [Majer and Peterson, 2007]. Injection rates increased significantly in 1997 with the start of the Southeast Geysers Effluent Recycling Project (SEGEP) (Figure 2.11). This project involves the delivery of treated wastewater and lake waters from Lake County to The Geysers geothermal field via a 46.4 km pipeline at a rate of about 22 million L/day [Majer et al., 2007]. Another reinjection project, the Santa Rosa Geothermal Reinjection Project (SRGRP), commenced in 2003 when a 64 km pipeline started sending treated wastewater from the city of Santa Rosa to The Geysers at a rate of about 30 million L/day (Figure 2.11) [Majer et al., 2007]. The depth of the recharge reservoir is between 2,134 and 3,048 m [Karner, 2005].

Figure 2.11 Historical seismicity from 1965 (start of significant production) to October 2006 at The Geysers. M represents the local magnitude. The largest recorded event had a magnitude 4.6 and occurred in 1982. Notice the increases in water injection in 1997 and 2003 due to the SEGEP and SRGRP, respectively [source: Majer and Peterson, 2007].

At present, The Geysers is one of the most seismically active areas in northern California [Mossop and Segall, unpublished study]. Mossop and Segall (unpublished study) found that the microseismicity at The Geysers is associated with both steam production and water injection during the period 1976-1998. They distinguished between three types of induced seismicity: i) Shallow seismicity (z ≤ 1 km) due to steam production with a time lag of more than 16 months; ii)





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Deep seismicity (z ≥ 1 km) due to water injection with a time lag ≤ 2 months and iii) Deep seismicity (z ≥ 1 km ) due to steam production with a time lag ≤ 2 months. The third type of induced seismicity becomes less significant after the mid 1980s [Mossop and Segall, unpublished study].

Information on historical seismicity (1965-2006) in relation to steam production and water injection is visualised in Figure 2.11. Since the late 1980s, the level of seismicity (of M ≥ 1.5) does not correlate with steam production, but with the amount of injected fluids [Majer et al., 2007]. The occurrence of LME events (M ≥ 4.0) show an increase with time and therefore increase with the amount of fluid injection (Figure 2.11) [Majer and Peterson, 2007]. Peaks in seismicity occurred in 1986 and in 1998 and recent increases in seismicity are due to the latest injections [Majer and Peterson, 2007]. Since the decrease in steam production (from 1986), the level of seismicity has remained fairly constant [Majer and Peterson, 2007]. The largest event with local magnitude 4.6 took place in the 1980s when steam production was at its peak [Majer et al., 2007]. Since that time, only a few events with magnitude 4 have occurred [Majer et al., 2007]. 4.2.1.10 Cooper Basin (Australia) The Cooper Basin geothermal site is located in South Australia and has a granitic reservoir at a depth of ~3.6 km overlain by sediments (Figure 2.12) [Baisch et al., 2006]. In 2003, the first well (Habanero-1) was drilled to a depth of 4,421 m where temperatures are approximately 250° C [Asanuma et al., 2004]. High overpressures exist in the granite joint network which resulted in the natural hydraulic stimulation over geothermal time [IEA, 2003]. Therefore the joints with favourable orientation (subhorizontal) were already permeable [IEA, 2003].

In November and December 2003, more than 20,000 m3 of water was injected in the well with maximum injection rates of 48 l s-1 and a maximum wellhead pressure of 65.5 MPa [IEA, 2003; Asanuma et al., 2004; Baisch et al., 2006]. About 27,000 induced events were detected during stimulation of the reservoir and were correlated to the amount of volume injected into the reservoir [Asanuma et al., 2004; Baisch et al., 2006]. The largest event (M=3.7) occurred after shut-in of the injection well [Baisch et al., 2010]. Since the Cooper Basin is a remote area, the occurrence of large events did not raise much community concern [Majer et al., 2007].

In contrast to the vertical seismic clouds at Soultz-sous-Forêts and Basel, the seismic cloud has a flat pancake shape (Figure 2.13) [IEA, 2003]. Seismicity started at the injection well and migrated away from it with time [Baisch et al., 2009]. The volume of the cloud is large in comparison to other EGS sites and is probably due to the large overpressures in the granite [IEA, 2003]. However, the number of events per volume of seismic cloud (seismic density) is much lower for the Cooper Basin than for Soultz-sous-Forêts [Asanuma et al., 2004]. As it is generally assumed that seismic density correlates





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Figure 2.12 The Cooper Basin geothermal site in Australia [Google Earth]. to the improvement of permeability after stimulation, this suggests that permeability has not increased as much as it has for Soultz-sous-Forêts [Asanuma et al., 2004]. This suggests that the fracture zone was not created during hydraulic stimulation, but that it was already present due to the tectonic history [Baisch et al., 2006]. The induced seismicity therefore occurred on the subparallel horizontal fractures which were already present [Asanuma et al., 2004].

Habanero-1 was restimulated in September 2005 by injecting 22,500 m3 of water for 13 days with a maximum wellhead pressure of 62 MPa [Baisch et al., 2009]. The goal of this stimulation was to enlarge the stimulated reservoir and to further enhance the permeability of the granite [Baisch et al., 2010]. Approximately 16,000 seismic events were induced during this injection period [Baisch et al., 2009]. The hypocenters of the events formed a flat subhorizontal structure resembling the shape observed during the first stimulation [Baisch et al., 2010]. The only difference is that the cloud had grown into the northern and western directions [Baisch et al., 2009]. Also, the spatiotemporal distribution of seismicity was different for this stimulation. Seismic activity started at the rim of the zone, which was previously stimulated, and then migrated towards the injection well and also into the opposite direction [Baisch et al., 2009]. In contrast to the 2003 stimulation during which seismicity started almost immediately, the time between start of injection and onset of induced seismicity was approximately 22 hours [Baisch et al., 2009]. The largest event had a local magnitude of 2.9 and also occurred after shut-in of the well [Baisch et al., 2010].





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Figure 2.13 Absolute hypocenter locations from the 2003 stimulation at Cooper Basin, in map view (top) and in perspective view along a northeast line (bottom). Origin time is denoted by the grey shading [source: Baisch et al., 2006].





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4.2.1.11 Berlín (El Salvador) The Berlín geothermal field - located on the flanks of the dormant volcano Cerro Tecapa - is one of six geothermal fields in El Salvador and was developed in the 1990s (Figure 2.14) [Bommer et al., 2006]. El Salvador is a region characterised by high seismic activity with most events generated by the subduction of the Cocos plate beneath the Caribbean plate [Bommer et al., 2006]. Shallower crustal events are due to the chain of Quaternary volcanoes [Bommer et al., 2006]. Local faults at the Berlín site are mostly oriented NW-SE [Ruggieri et al., 2006]. The Berlín field consists of eight production wells and a reinjection system, which consists of ten injection wells [Bommer et al., 2006]. The depths of these wells range between 700 m and 2,500 m [Bommer et al., 2006]. The reservoir at the Berlín field consists of basaltic-andesite and andesite lavas interbedded with minor tuff layers [Raymond et al., 2005]. The maximum temperature measured in the wells is up to 305° C [Tassi et al., 2007].

Figure 2.14 Map showing the locations of the geothermal sites at Berlín (El Salvador) and Bouillante (French Antilles) [Google Earth].

A large hydraulic stimulation, involving three injections, took place in an inclined well

(TR8-A), which extends to a TVD of about 2,466 m [Oates et al., 2004]. During the first injection -





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which took place between 28 June and 17 July 2003 - the shallower formations at around 1,800 m MD were stimulated [Oates et al., 2004]. Between the 18th of August and the 2nd of September the shallower formations were stimulated again [Oates et al., 2004]. The deeper reservoir interval was stimulated during the third hydraulic injection, which took place between 22 December 2003 and 8 January 2004 [Oates et al., 2004].

A total of 300,000 m3 was injected during the entire stimulation period [figure 1 in Bommer et al., 2006]. The maximum injection rate during the first stimulation period was ~108 l s-

1 with a maximum wellhead pressure of ~7.8 MPa [figure 12.1 in Oates et al., 2004]. The second injection took place at maximum injection rates of ~145 l s-1 and maximum wellhead pressure of ~9 MPa [figure 12.2 in Oates et al., 2004]. A maximum injection rate of 100 l s-1 and a maximum wellhead pressure of 13.5 MPa were used for the third injection period [figure 12.3 in Oates et al., 2004].

Microseismicity during the entire stimulation period was very low; 32 events, with magnitude greater than or equal to 0.5, occurred during the first injection test, 49 during the second injection test and during the final injection period 24 events occurred [Oates et al., 2004]. During periods of injection the event rate was doubled from 1.1 events/day to 2.4 events/day [Oates et al., 2004]. The largest event had a local magnitude of 4.4 and took place on 16 September 2003, two weeks after shut-in of the second injection period [Bommer et al., 2006]. Hypocenter locations of the induced events during these hydraulic injections are shown in Figure 2.15. The timing and location (~ 3 km to the south of TR8-A) of this event raised questions about the origin (induced/tectonic) of this LME event [Bommer et al., 2006]. At other injection sites, LME’s also occurred after stimulation (e.g. at Soultz-Sous-Forêts, Basel and Cooper Basin) supporting the idea that the event was triggered by the injection [Bommer et al., 2006].

Seismicity was much lower than would be expected compared to the stimulations at the Soultz-sous-Forêts site, even though injection rates were much higher for the Berlín stimulation [Oates et al., 2004]. Oates et al. (2004) argue that this cannot be explained by the difference in strength between the volcanic rocks and the granite. After they corrected for the weaker volcanic rocks, they still did not expect the low level of seismicity observed at Berlín. Therefore, Oates et al. (2004) suggest that the simplest explanation for the difference in seismicity is that large-scale hydraulic stimulation - as observed at Soultz-Sous-Forêts - did not take place in Berlín.

4.2.1.12 Bouillante (French Antilles) The Bouillante geothermal field is located on the Basse-Terne Island which is part of the Guadeloupe archipelago (Lesser Antilles) (Figure 2.14) [Traineau et al., 1997]. This geothermal field is located where the Atlantic lithosphere is subducting southwestward beneath the Caribbean plate [Bouchot et al., 2010]. The geothermal system consists of submarine volcanoclastic formations (hyaloclastites, scarce lava flows) and subaerial formations (andesitic lava flows, pyroclastites, lahars) [Mas et al., 2006]. The top of the heat reservoir becomes deeper going from the north (~300 m depth) to the





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Figure 2.15 Vertical cross-sections showing the hypocenter locations of induced events during periods of hydraulic stimulation at Berlín (June 2003-January 2004). The bottom figure shows the same locations as the top figure, but the bottom figure is viewed from the North of TR8-A [source: Oates et al., 2004].





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south (~600 m) and is located at the base of an illite smectite alteration zone [Bouchot et al., 2010]. This high enthalpy geothermal system has temperatures of 240° C at 350 m depth [Traineau et al., 1997]. The geothermal activity at Bouillante could be related to fissural volcanic activity, which is probably related to the Basse-Terre-Montserrat fault system which is oriented NNW-SSE [Mas et al., 2006].

Small-scale near-vertical joints are distributed over the area [Traineau et al., 1997]. These joints can be grouped into a major set having a strike of N110° and a minor set with a strike of N010° [Traineau et al., 1997]. Normal faults are also present and are mainly striking N100-120° and dipping 90 to 70° [Traineau et al., 1997]. These normal faults have an offset of around a few meters [Traineau et al., 1997]. Steeply dipping filled fractures (extensional) are also present in this area and can be subdivided into a major and a minor set with a strike of N110-120° and N080-090°, respectively [Traineau et al., 1997]. The anisotropic distribution of these fractures probably resulted in the large contrasts in permeability throughout the reservoir[Traineau et al., 1997].

In the 1970s the first four exploratory wells (BO-1, BO-2, BO-3 and BO-4) were drilled in the fractured reservoir to depths ranging between 350 and 2,500 m depth [Traineau et al., 1997]. Well BO-2 (350 m depth) was the only well out of these four which showed significant steam production, due to the intersection of a permeable fracture network [Traineau et al., 1997]. Between 1986 and 1992 and between 1996 and 2002, BO-2 was producing about 150 t/h of fluid (30 t/h steam), exploiting the top part of the reservoir [Mas et al., 2006; Lopez et al., 2010]. Well BO-4, which was drilled to a depth of 2,500 m, showed less productivity (15 t/h steam), whereas well BO-1 and B0-2 were abandoned [Traineau et al., 1997].

In 2001, three directional wells were drilled: BO-5, BO-6 and BO-7 [Mas et al., 2006]. Well BO-5 and BO-6 were the only productive wells out of these three wells [Mas et al., 2006]. After well BO-2 was shut in July 2002, wells BO-5 and BO-6 started alternatively producing fluid [Lopez et al., 2010]. No produced fluid is reinjected, because the reservoir is recharged by sea water and fresh surface water [Bouchot et al., 2010; Sanjuan et al., 2010a].

Estimates of annually fluid production are 4.5 millions of tons and pressure has dropped from 1.3 to about 0.9 MPa since the increase (from 150 to 600 tons/h) in production in 2005 [Lopez et al., 2010]. Tectonic earthquakes have been recorded at the Bouillante field, indicating that this area is seismically active [Sanjuan et al., 2010a]. However, production-induced seismicity was not recorded at Bouillante, suggesting that fluid production does not result in short-term pressure changes of the reservoir [Sanjuan et al., 2010a]. 4.2.2 Sites not included in the GEISER project

4.2.2.1 Fenton Hill (New Mexico, US) The Fenton Hill geothermal site is located in New Mexico (US) and has a reservoir consisting of granitic rocks (Figure 2.10) [Dinske, 2010]. This geothermal site is situated on the western boundary of the Rio Grande rift, which is active today and characterised by high seismicity [Barton et al., 1988].

The development of a deep reservoir began with the drilling of well EE-2 to a vertical depth of about 4,400 m with a bore hole temperature of 323° C [Laughlin et al., 1983; Brown and Duchane, 1999]. The lower part of this well was drilled toward the east at an angle of 35° from the vertical [Brown and Duchane, 1999]. A second well (EE-3) was drilled to 3,980 m TVD in a similar way, positioned 46 m at the surface from EE-2 [Laughlin et al., 1983; Brown and Duchane, 1999]. A massive hydraulic fracturing experiment took place in December 1983 by the injection of over 21,000 m3 of water into EE-2 [Dinske, 2010]. This stimulation lasted for about 62 hours with average injection rates of 100 l s-1 and a maximum wellhead pressure of ~48 MPa [Shapiro, 2008; Dinske, 2010]. About 11,366 event were recorded, from which 9,350 events were recorded during injection [Shapiro, 2008]. The occurrence of microseismicity after injection was also





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observed at other injection sites such as Soultz-sous-Forêts [Dinske, 2010]. Most of these events were small magnitude events, indicated by a high b-value in the Gutenberg-Richter plot [Philips et al., 2002]. The hypocenter locations form an elongated zone with a strike of N355° and a dip of E70°, very similar to the form of the seismic cloud at the 1993 injection at Soultz-sous-Forêts [Philips et al., 2002; Dinske, 2010]. The orientation of the seismic volume suggests that the major fracture set, which was opened due to stimulation, had about the same inclination as the drilled well [Brown and Duchane, 1999]. This means that the wells EE-2 and EE-3 were drilled unfavourably for EGS, just as the RH11 and RH12 wells at the Rosemanowes site [Brown and Duchane, 1999]. In May 1984, a hydraulic injection took place in well EE-3 [Fehler, 1989]. During this stimulation, 7,600 m3 of water was injected with an average rate of 20 l s-1 [Fehler, 1989]. Again, microseismicity developed along the direction of the well [Brown and Duchane, 1999]. After EE-3 was deviated from its original trajectory, more injection experiments were carried out in this new well (EE-3a) having a vertical depth of 4,018 m and a bottom hole temperature of 265° C [Fehler, 1989; Brown and Duchane, 1999; Tester et al., 2006]. In June and July of 1985, EE-3a was stimulated by injection of 5,600 m3 of water at an average rate of 20 l s-1 [Fehler, 1989]. Another stimulation of EE-3a took place in January and February of 1986 when 4,400 m3 was injected at an average rate of 20 l s-1, but at slightly shallower depth than in 1985 [Fehler, 1989]. The stimulations in 1984 and 1985 induced more microearthquakes than the stimulations in 1983 and 1986 [Fehler, 1989]. During the 1984 and 1985 stimulations, microseismicity was distributed along one or two planes [Fehler, 1989]. The other two stimulations, on the other hand, resulted in microseismicity which was more evenly distributed along more planes [Fehler, 1989]. Fehler (1989) suggests that planes, which are favourably oriented for slip, might be present in the proximity of the injection points in the 1984 and 1985 experiments. This may have led to the higher amount of induced events during these two experiments [Fehler, 1989]. 4.2.2.2 Ogachi (Japan) The Ogachi geothermal site is located in the northeast of Japan near Kurikoma National Park on Honshu Island (Figure 2.16) [Tester et al., 2006]. This geothermal site is situated in a caldera and has granodiorite reservoir rocks [Kaieda et al., 2010]. Natural joints in the reservoir rock are predominantly north-north-east oriented [Kaieda et al., 2010]. In 1990, well OGC-1 was drilled to a depth of 1,000 m with a bottom hole temperature of over 230º C [Tester et al., 2006]. After fracture initiation tests, OGC-1 was stimulated in 1991 by injecting over 10,000 m3 of water into its open-hole section [Shapiro et al., 2007; Kaieda et al., 2010]. The purpose of this injection was to show that large reservoirs can be created even at shallow depth [Kaieda et al., 2010]. Injection took place during eleven days with maximum injection rates of ~12.5 L s-1 and maximum wellhead pressure of ~19 MPa [figure 1 in Shapiro et al., 2007].

During the first five hours of water injection, seismicity was low and only a few events were observed [Kaieda et al., 2010]. Kaieda et al. (2010) suggest that this is due to the Kaizer effect; no events occur until the injected volume becomes larger than the previous injected volume. The largest event had a magnitude 2.0, while the magnitude of all the other events was below -1.0 [Kaieda et al., 2010]. Kaieda et al. (2010) explain the low magnitude events by the occurrence of a large amount of joints and relatively low stress conditions at Ogachi [Kaieda et al., 2010]. The direction of hypocenters of the induced events was oriented along the strike of the natural joints (NNE) [Kaieda et al., 2010].





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Figure 2.16 Map showing the location of the geothermal site at Ogachi (Japan) [Google Earth].

At Ogachi and Cooper Basin very similar microearthquake observation systems were used, therefore seismicity at these two geothermal sites can be compared [Kaieda et al., 2010]. In contrast to Ogachi, seismic activity was very high during stimulation at Cooper Basin and commenced very quickly after start of injection. It is suggested by Kaieda et al. (2010) that seismicity at Ogachi did not start immediately after the start of injection, because much of the injected water was recovered to the surface resulting in relaxation of the opened fractures. The water injected before the start of the main stimulation at Cooper Basin was not allowed to flow back to the surface [Kaieda et al., 2010]. Therefore reservoir pressure remained high and resulted in the observed initial high seismicity [Kaieda et al., 2010]. Another difference between these two sites is the volume of rock which was seismically activated [Kaieda et al., 2010]. This volume is much larger for the Cooper Basin stimulation, which Kaieda et al. (2010) explain by the difference in the natural joint system at both sites. In contrast to Ogachi, only few natural joints are present at Cooper Basin and stress conditions are high [Kaieda et al., 2010]. Kaieda et al. (2010) suggest that at Cooper Basin a small number of natural joints was stimulated concentrically leading to large magnitude events and the stimulation of a large area. 4.2.2.3 Cotton Valley (Texas, US) The Cotton Valley formation (Texas) - consisting of low permeable gas-bearing sands - has been hydraulically stimulated to increase the production of this gas field (Figure 2.10) [Rutledge et al., 2004; Fischer and Guest, 2011]. Natural fractures present in the Cotton Valley sand intervals are mostly vertical extension fractures, oriented subparallel (within 10°) to the direction of the maximum horizontal stress (80°) [Fischer and Guest, 2011; Rutledge and Phillips,





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2003]. The shales which are present in between the sandstone layers only contain few natural fractures [Rutledge and Phillips, 2003]. Reservoir temperatures are about 120° C [Dinske, 2010]. Gel-proppant treatments (A and B) were performed in well 21-10 (~2,980 m depth) in May 1997, whereas waterfrac treatments were conducted in the adjacent well 21-09 (~2,780 m depth) in July 1997 [Rutledge et al., 2001; Rutledge et al., 2004; Dinske, 2010]. Treatment A involved the injection of 1,340 m3 of slurry, containing 228,634 kg proppant, at a rate of 119 l s-1 into a perforated interval between 2,615 and 2,696 m depth [Rutledge et al., 2004]. 1,122 induced events were recorded during this treatment [Rutledge et al., 2001]. Injection of 1,253 m3 of slurry with 190,690 kg proppant took place during the other gel proppant treatment in well 21-10 [Rutledge et al., 2004]. This stimulation took place at a deeper depth interval with injection rates of 106 l s-1 and nearly 1,200 events were detected [Phillips et al., 2002; Rutledge et al., 2004].

During treatment C, 419 m3 of slurry - containing 14,891 kg of proppant - was injected in the perforated depth interval between 2,607 and 2,643 m of well 21-09. Injection rates for the stimulations in this well were much lower than those in well 21-10; a rate of 26.5 l s-1 was used for both treatment C and D and rates between 21.2 and 26.5 l s-1 were used during treatment E [Rutledge et al., 2004]. Treatment D involved the injection of 396 m3 of slurry containing 12,394 kg of proppant in the interval between 2,663 and 2,687 m depth [Rutledge et al., 2004]. 400 m3 of slurry with 7,264 kg of proppant was injected in the depth interval between 2,746 and 2,763 m during treatment E [Rutledge et al., 2004].

The hypocenters of the induced earthquakes lie along narrow bands of seismicity trending along vertical fractures which are oriented within 5º to the maximum horizontal stress [Fischer and Guest, 2011]. For treatments B and E (deeper depth interval), seismicity also occurred along fracture sets oriented 20-30º to the maximum horizontal stress [Fischer and Guest, 2011]. The narrow bands of seismicity are contained to the sandstone layers which were stimulated, this is shown in Figure 2.17 for the induced





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Figure 2.17 Vertical cross-section through wells 21-10 and 21-09 at Cotton Valley. Hypocenter locations are shown for treatment A (well 21-10) [source: Rutledge and Phillips, 2003]. seismicity during treatment A [Rutledge et al., 2004]. This suggests that the sand intervals are hydraulically isolated [Rutledge and Phillips, 2001]. This also indicates that the natural fractures present in the reservoir were stimulated, because they are mostly present in the sand intervals [Rutledge and Phillips, 2003]. The Cotton Valley reservoir is characterised by high b-values, which means that large events are less likely to occur [Dinske, 2010]. This is due to the smaller amount of volume injected, since injection was not aimed at developing a geothermal reservoir [Dinske, 2010].

Phillips et al. (2002) studied the pattern of hypocenter locations induced during hydraulic stimulations of sedimentary, limestone and crystalline reservoir. They observed that for clastic sedimentary reservoirs microseismicity is characterised by horizontal, linear patterns. This is evident for the hydraulic stimulations at the Cotton Valley gas field. They also showed that microseismicity in crystalline rocks and massive, brittle sediments show combinations of linear and planar features, indicating interacting networks of fractures.





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4.2.2.4 Barnett Shale (Texas, US)

Gas is produced from the Barnett Shale formation located in the Fort Worth Basin, Texas (Figure 2.10) [Dinske, 2010]. The Barnett Shale gas play is the largest gas field in Texas [Gale et al., 2007]. However, this gas-bearing shale rock is extremely impermeable and therefore hydraulic fracturing is needed to produce gas from the Barnett Shale Formation [Dinske, 2010]. In Fort Worth Basin, at least two sets of natural fractures are present; an older north-south-trending set and a dominant, younger west-northwest-east-southeast trending set [Gale et al., 2007]. The natural fractures, present in the Barnett Shale, are mostly subvertical [Gale et al., 2007]. Orientation of maximum horizontal stress is NE-SW in the Fort Worth Basin, directed normal to the orientation of dominant natural fractures [Gale et al., 2007]. Hydraulic fractures propagate in the direction of the maximum horizontal stress, until they come across natural fractures. Therefore, the wells are commonly drilled normal to the maximum horizontal stress in order to maximise the stimulated volume [Gale et al., 2007].

In the Summer of 2001, such a fracture treatment was conducted by injection of 2,840 m3 of water at high injection rates of 150 l s-1 and wellhead pressures between 40 and 42 MPa [Dinske, 2010]. This stimulation lasted for 5.5 hours and the non-planar growth of microseismicity indicates that the injected fluids opened up a complex pre-existing fracture network [Dinske, 2010]. 4.2.2.5 Paradox Valley (Colorado, US) The Dolores River, a tributary of the Colorado river, adds more than 2·108 kg total dissolved salts/year to the Colorado River’s salinity [Mahrer et al., 2005]. The Dolores River receives these large amounts of salt by seepage from a briny aquifer on its way through Paradox Valley, a NW trending collapsed diapiric salt anticline [Mahrer et al., 2005]. The Paradox Valley (Southwestern Colorado) is 40 km long and 8 km wide and is underlain by about 6 km of interbedded salts and shales (Figure 2.10) [Mahrer et al., 2005].

The purpose of the Paradox Valley unit (PVU), a U.S. Bureau of Reclamation Project, is to reduce the contribution of the Dolores River to the Colorado River’s salinity and therefore improving the quality of the Colorado River [Mahrer et al., 2005; Cladouhos et al., 2010]. The study of this brine seepage problem commenced in 1971 and in the late 1970s about 90 shallow (12-21 m) extraction, monitoring and testing wells, along the Dolores River in Paradox Valley, were in operation by the PVU [Mahrer et al., 2005]. At present, the PVU uses 9 of these wells to extracts the brine from the aquifer before it enters the Dolores and injects it into a 4.9 km deep injection well (No. 1) after treating the water [Mahrer et al., 2005].

The diluted brine is injected into the Leadville formation, a highly fractured dolomitic limestone (Figure 2.18) [Mahrer et al., 2005]. The well is perforated between 4.3 and 4.9 km, including ~200 m of Leadville and ~70 m of aphanitic schist in the perforated interval [Mahrer et al., 2005]. The location of the injection well was planned





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Figure 2.18 Cross-section normal to the northwest axis of Paradox Valley, containing the location of the PVU injection well and the implied faults of the Wray Mesa fault system. Locations of injection-induced events during 1991-2002 are also shown [source: Mahrer et al., 2005]. so that it would optimise fluid migration into and along the faults of the Wray Mesa Fault System, present in the Leadville and surrounding formations (Figure 2.18) [Mahrer et al., 2005; Cladouhos et al., 2010]. These faults dip 80º to the NE and the primary faults of the Wray Mesa system strike N55º W, subparallel to the long axis of the Paradox Valley [Mahrer et al., 2005].

Due to the injections at Paradox Valley, seismicity has been monitored over three time periods: pre-injection (1985-1991); injection testing (1991-1995) and continuous injection (1996-present) [Mahrer et al., 2005]. In general, Paradox Valley is considered to be a seismically inactive region; between 1985-1991 only six minor events were detected and none of them were located within 10 km of the injection well [Mahrer et al., 2005]. Seven injection tests were performed between July 1991 and April 1995 [Mahrer et al., 2005]. In these injection tests waste brine was continuously injected with a wellhead pressure between 25-32 MPa and each test was followed by a wellhead shut-in [Mahrer et al., 2005]. Injection rates varied between 9 and 25 l s-1 and the in-situ pressure before injection was 43.6 MPa [Cladouhos et al., 2010]. In total 620,072 m3 was injected and 666 events were recorded during 438 days of pumping [Mahrer et al., 2008]. The first event started three days after the beginning of the first injection test [Mahrer et al., 2008]. Continuous injection started in May 1996 [Mahrer et al., 2008]. Between May 1996 and October 2004 about 3,435 induced events were located [Mahrer et al., 2005]. So a total of more than 4,100 events were recorded at the surface, during 1991-2004 [Mahrer et al., 2005]. Over 99.9 % of these events was below magnitude 2.0, so below the human detection level [Mahrer et al., 2005]. Since the start of continuous injection, a dozen events occurred which were felt by humans, with the first event taking place in August 1997 [Mahrer et al., 2005]. The continuous injection period can be subdivided into four phases: Phase I through Phase IV [Mahrer et al., 2005]. These phases correspond to changes in injection parameters to mitigate induced seismicity [Ake et al., 2005; Mahrer et al., 2005].





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Phase I started in May 1996 with the injection of fluid at a rate of 21.5 l s-1, a surface pressure of ~33 MPa and a down hole pressure of ~80 MPa [Mahrer et al., 2005]. In May 1999, a magnitude 3.5 event occurred [Mahrer et al., 2005].

Phase II started after the occurrence of a magnitude 3.6 event in June 1999 [Mahrer et al., 2005]. During this phase, the same injection pressure and injection rate were used as in Phase I, but a 20-day shutdown was introduced every six months [Mahrer et al., 2005]. This measure resulted in the reduction of induced seismicity, but it did not prevent large induced events from happening [Mahrer et al., 2005]. The largest event occurred in May 2000, having a magnitude 4.3 [Mahrer et al., 2005].

Phase III followed after this event and during this period injection rates were lowered to 14.5 l s-1, bottom hole pressure stayed the same whereas surface pressure was reduced by about 10 % [Mahrer et al., 2005]. Both the 20-day shut-ins and the lowering of the injection rates helped in reducing the amount of induced events [Mahrer et al., 2005]. Between 1998 and the occurrence of the magnitude 4.3 event the average number of events per month was 81, while between late June 2000 and the end of 2002 this was only 9 events/month [Mahrer et al., 2005]. Since then it has been reduced even more to 5 events/month in 2002 [Mahrer et al., 2005].

In January 2002 Phase IV commenced, in this period the same injection rate and pressure as in Phase III was used with a 20-day shut-in every six months [Mahrer et al., 2005]. At the beginning of Phase IV, surface pressure was about 30 MPa with a down hole pressure of 79 MPa [Mahrer et al., 2005]. In November 2003 both the surface pressure and down hole pressure had increased to 32 MPa and 81 MPa, respectively [Mahrer et al., 2005].

The hypocenters of the seismic events were all located in a narrow depth range, following the depth interval of the Leadville formation by dipping 15º to the east (Figure 2.18) [Mahrer et al., 2005]. The hypocenters can be divided into two separate seismic clouds, one large group of events located asymmetrically around the injection well and a smaller cloud along the trend of the Wray Mesa fault system situated about 8 km to the northwest [Mahrer et al., 2005]. The induced events all occurred on pre-existing faults, joints and planes of weakness [Mahrer et al., 2005].

4.3. Key parameter identification The parameters in the database were selected based on prior knowledge of gas depletion induced seismicity and existing research for EGS induced seismicity. There are parameters for site factors (such as geology of the site and in-situ stress), operational parameters (such as injection rates and pressures) and parameters for the observed induced seismicity (such as maximum observed magnitude and number of events). The values of the parameters for these different injection cases are described in Appendix I. A first selection of these parameters with potential for correlation with the observed maximum magnitude at an injection site is taken from this database. This first selection was based on work by Shapiro (2008). Shapiro modelled the induced seismicity and obtained good comparisons with field cases. The triggering front of seismicity (Shapiro, 2008) is related to the time of injection and hydraulic diffusivity. The source strength (sterength of the induced event) (Shapiro, 2008) is related to the bottom hole pressure, radius of the borehole and hydraulic diffusivity. Shapiro (2008) also identified the tectonic potential of the medium which relates the criticality of the medium and the density of the pre-existing cracks. Other parameters are permeability, minimum horizontal stress, pore pressure perturbation and critical pore pressure. The observed maximum magnitude not the total cumulative energy is chosen for the comparison, contrary to the study by Van Eijs et al. (2006), is chosen as this can easily be found in the literature and of imminent interest to the public. Although both seismic energy as well as Mmax should give similar results, there are indications that the cumulative seismic energy versus cumulative volume of injection might be important (Thomas Kohl, personal communications). The cumulative seismic energy is difficult to determine due to the lack of seismic data for most sites (except perhaps Soulz-sous-Forêts).





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Figure 3.1 shows the following selected parameters: volume of seismic cloud (Vseism); injected volume (Vinj); maximum formation pressure (Pformation); maximum wellhead pressure (Pwh); minimum horizontal stress (Sh); maximum injection rate (Qmax) and duration of injection (Tinj). These parameters were made dimensionless to remove any effects of specific dimensions of individual parameters. The dimensionless parameters chosen here are Vseism/Vinj, Pformation/Pwh, Sh/Pwh and Qmax*(Tinj/Vinj). The volume of the seismic cloud will depend on the volume which is injected. The injection rate is normalised by the duration of the injection over the injected volume, which gives a measure of the relative injection rates at the different sites and for different injections. We have also plotted some non-dimensionless parameters (Appendix III).

In Figure 3.1a)-d), the maximum observed local magnitude is given as a function of the dimensionaless parameters for several injection sites. Overall, there is no correlation visible in the four graphs. Bayesian least squares fitting was performed on the data points using logarithmic (Mmax=a*xb) and linear functions (Mmax= a+bx), with x the chosen dimensionless parameter. Bayesian least squares starts by multiplying the probabilities for the observations (Gaussians) with the distribution (assumed Gaussian) for the parameters of the fit as they are expected a priori. The uncertainty for the observed Mmax is taken as a priori information in this analysis and is set to 0.16. The value of 0.16 was derived from differences in reported magnitudes for registrations of the same events in worldwide catalogues. For all of these fits large values for Χ2 and the point average were found, Χ2 represents the sum of the difference between observed Oi and estimated (from fit) values Ei divided by the uncertainties for the data:

ii

k

ii EO σ/)( 2

1

2 −=Χ ∑=

2 (1)

and the average X2 is X2 divided by the number of data points. So this means that these are not good fits to the data. An example of a logarithmic fit through the data points from Figure 3.1c) is shown in Figure 3.2.

In Figure 3.1a), the observed maximum local magnitude is plotted as a function of Vseism/Vinj . The data in this graph suggests that mainly larger events (≥2.0) occur in granitic reservoir rock and that smaller events occur in other reservoir rock (e.g. basalt, limestone, gneisses and limestone). An exception might be the magnitude 4.4 event at Berlín (basalt/andesite reservoir rock), but there are indications that this is a tectonic event instead of an injection-induced event.





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Plotting of Pformation/Pwh shows that all the injections into granite have a magnitude exceeding 2.0 (Figure 3.1b)). However, three data points, belonging to Paradox Valley (injection into limestone), have a magnitude above 2.0. The high magnitudes at Paradox Valley could be due to the Wray Mesa Fault system, which is intersected by the deep injection well at this site. Figure 3.1c), in which the observed maximum local magnitude is plotted as a function of Sh/Pwh, again shows the division between high magnitude events in granite rocks and low magnitude events in other reservoir rocks. This is also the case for Figure 3.1d) in which Qmax*(Tinj/Vinj) is plotted (with the exception of the 4.4 event at Berlín).

Figure 3.1 Maximum local magnitude plotted as a function of a) Vseism/Vinj; b) Pformation/Pwh; c) Sh/Pwh and d) Qmax*(Tinj/Vinj) (from database in Appendix I). Sites are separated into those which have a granitic reservoir (blue) and those which have other reservoir rock (green). The dashed line at Mmax of 2.0 represents the assumed division between seismic events which are felt and those which are not felt by humans. The outlier in a) belongs to the injection at KTB in 1994, where a relatively small amount of water was injected in comparison to the seismic volume. The three green points in b) with M> 2.0 belong to the injection site at Paradox Valley, where the injection well was intersected by the Wray Mesa fault system.

a) b)

c) d)





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Figure 3.2 Bayesian least squares fitting (with an uncertainty of 0.16 for Mmax) gives Χ2 = 105.891 and average Χ2= 11.766 (with an uncertainty of 0.16 for Mmax). The point for Groβ Schönebeck has been excluded from this logarithmic fit, because of the negative Mmax for this site. The formula for the logarithmic fit is also given. 4.4. Discussion and conclusions

There are only few data points in the Figures 3.1a)-d) and Bayesian least squares fitting of the data resulted in bad fits (large values for Χ2 and the point average). The bad Bayesian fit suggests the wrong underlying assumptions, i.e. that the dimensionless parameters have a relation with the maximum observed magnitude. The values for Χ2 and the point average are mainly large, because the uncertainties for Mmax are significantly smaller than the deviations from the data points to the fit. The small amount of data points is mainly due to the lack of reported values on the proposed dimensionless parameters and the observed maximum magnitude (Mmax). First of all, only a small amount of injection sites has been included in this study (19). Also, Mmax could not be found for sites which have not been extensively described in the literature. This means that these sites had to be taken out from the data represented in Figures 3.1a)-d), since the parameters are plotted as a function of Mmax. These sites include Rosemanowes, Fenton Hill, Cotton Valley and the Barnett Shale at Fort Worth Basin. Also, the amount of numerical data on parameters differed for each of the sites. This means that if a certain parameter is only found for a couple of sites, it cannot be used in the analysis.

Extensive information is available on some specific sites. These include Soultz-Sous-Forêts, Basel, Groβ Schönebeck, Cooper Basin and Paradox Valley. At The Geysers, continuous injection and production have been ongoing for years now using a large amount of injection and production wells, which makes this site more complicated. Deep drilling programs at Iceland and Campi Flegrei still need to be completed, therefore these sites also cannot be used in the analysis.

In the Van Eijs et al. study (2006), on induced seismicity related to gas depletion, gas fields were divided in classes on the basis of the occurrence of seismicity. This division cannot be applied for the injection sites in this paper, since seismicity always occurs even though at some sites induced seismicity is extremely low. For the induced events caused by gas depletion in the study by van Eijs et al. (2006), smaller events (M < 1.0) could have taken place but were not recorded due to the detection threshold of M > 1.5 for the whole of the Netherlands. A test with a borehole seismometer in the Roswinkel gas field, during the depletion phase, recorded 40 small events with magnitudes smaller than

Mmax = 3.433* (Sh/Pwh)^(-0.197)





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zero (M < 0), during two weeks. The detection threshold for EGS is usually much lower (M < 0), due to the borehole seismometers in place. A conclusion drawn from this is that induced events are likely to occur both with gas depletion as well as with EGS. The occurrence of felt events (assuming a threshold of human detection of magnitudes larger than 2.0) depends on the site and possibly on the cause of induced seismicity.

For EGS, injection into granite has a high likelihood of generating felt events. However, there are more sites which have granitic reservoir rock compared to sites with other reservoir rock which may introduce a bias. Conclusions about the type of reservoir rock considering LME events cannot be drawn as yet due to the small amount of data (19 sites) in this study. The injections into the volcanic and sandstone intervals at Groβ Schönebeck also illustrate that the type of reservoir rock might be very important for the amount and the level of induced seismicity. At this geothermal site only a small amount of events were induced with a low Mmax (in comparison to other injection sites), although very high injection rates were used. The addition of more sites could help to assess whether site-specific factors (e.g. type of reservoir rock) do play a role in the induction of large magnitude events. The study by Evans et al. (in press) showed similar observations; they found that the seismogenic response of injection into sedimentary rocks were overall lower than into crystalline rocks. However, the majority of the sedimentary sites involve balanced circulations whereas most data from crystalline rocks are from injection experiments. Therefore they could not draw solid conclusions on the seismogenic response at sites which do not have crystalline reservoir rock.

Based only on the available data in this study, no conclusions can be drawn on injection-induced seismicity. This problem may be solved by extending the data set with other sites (including non-geothermal sites like CO2 injection, gas injection, hydraulic fracturing, shale gas production and water injection). This research can be more challenging than known beforehand. The sites considered are very different from each other (different geology, operations, seismicity), which makes using statistics difficult. Merging of this dataset with the dataset created in the study by Evans et al. (in press) is already in progress. Performing a Bayesian least squares analysis on this extended merged dataset should give a clearer view on whether there is a correlation between certain key parameters and induced seismicity and whether such a relation may eventually be found. Using this extended merged database together with the expected Mmax (instead of the observed Mmax) might result in good correlations between global dimensionless parameters (which differ per site) and Mmax. However, if a relation between Mmax and one or several parameters (f(Mmax, par1, par2, par3, …) cannot be found, you might have to conclude that the induced seismicity very much depends on site-specific factors.

In the study by Van Eijs et al. (2006), all the data came from gas fields (> 200) in the Netherlands having similar reservoir rock. However, the sites used in this paper are distributed around the world and thus all have very different geologic settings. Therefore, it is possible that the seismic hazard at a site prior to injection cannot be assessed by doing a global statistical analysis and that the seismogenic response to injection is site-specific. In order to confirm this hypothesis, more sites in similar geological circumstances would need to be investigated. Also, it could be very useful to collect more extensive information on the existence of faults at the sites and the distance at which these faults are located from injection wells. Maximum magnitude events, occurring after shut-in, might have been triggered on faults nearby the injection well (Table 1.1). The 2003 injection at Soultz-sous-Forêts and the stimulation at Basel give proof that the existence of faults can result in the triggering of large magnitude events on faults after shut-in. These triggered events are likely to have larger magnitudes then ‘induced’ events, because a larger area can slip on the fault. This supports the idea that site-specific factors might play an important role in inducing LME’s. The dependence of the magnitude of events with depth is still to be investigated. TNO is currently investigating this topic using seismic data from the Netherlands. The located seismic events are listed in the KNMI catalogue. From a statistical point of view smaller events take place at shallower depth. This suggests that since EGS occurs at small depths, statistically EGS will induce small events. An anomalous case is the KTB site, where the two injections occurred at ~9 km depth into the KTB-HB well, which was intersected by large





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faults. The maximum induced seismicity was quite small (Ml=1.2 and Ml=0.7), especially compared to the Basel and Soultz-sous-Forêts site, where stimulation took place at ~5 km depth.

A priori information on injection-induced seismicity before production of a site is a challenging subject. At this moment only numerical modelling of the injection sites may give an indication of the seismic hazard at potential injection sites, provided enough geological information is available (elasticity parameters, faulting, permeability, porosity, geological formations, etc.) and enough understanding of the physical mechanisms. Due to the large number of triggered events on nearby faults, it might be advisable to stay away from faults at future EGS sites until the mechanism of induced seismicity is better understood. This severely limits the amount of sites, because the presence of faults usually also indicates higher temperatures, which makes these sites more interesting for EGS. Another venue is to regard the induced seismicity as an acceptable risk (for example in The Geysers, USA), which allows these sites to be used for EGS. After the drilling and testing phase more indications of future seismic activity may be obtained by other statistical methods or numerical methods (results of the GEISER project).





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4.5 References Ake, J., K. Mahrer, D. O’Connell, and L. Block (2005), Deep-injection and closely

monitored induced seismicity at Paradox Valley, Colorado, B. Seismol. Soc. Am., 95 (2), 664-683. Aki, K., M. Fehler, R. L. Aamodt, J.N. Albright, R. M. Potter, C. M. Pearson, and J. W.

Tester (1992), Interpretation of seismic data from hydraulic fracturing experiments at the Fenton Hill, New Mexico, Hot Dry Rock geothermal site, J. Geophys. Res., 87 (B2), 936-944.

Ármannsson, H., J. Benjamínsson, and A. W. A. Jeffrey (1989), Gas changes in the Krafla geothermal system, Iceland, Chem. Geol., 76, 175-196.

Asanuma, H., Y. Kumano, T. Izumi, N. Soma, H. Kaieda, Y. Aoyagi, K. Tezuka, D. Wyborn, and H. Niitsuma (2004), Microseismic monitoring of a stimulation of HDR reservoir at Cooper Basin, Australia, by the Japanese team, Geoth. Res. T., 28, 191-195.

Asanuma, H., Y. Kumano, A. Hotta, U. Schanz, H. Niitsuma, and M. Häring (2007), Analysis of microseismic events from a stimulation at Basel, Switzerland, Geoth. Res. T., 31, 265-270.

Baer, M., N. Deichmann, J. Braunmiller, J. Clinton, S. Husen, D. Fäh, D. Giardini, P. Kästli, U. Kradolfer, and S. Wiemer (2007), Earthquakes in Switzerland and surrounding regions during 2006, Swiss J. Geosci., 100, 517-528.

Baisch, S., M. Bohnhoff, L. Ceranna, Y. Tu, and H.-P. Harjes (2002), Probing the crust to 9-km depth: Fluid-injection experiments and induced seismicity at the KTB superdeep drilling hole, Germany, B. Seismol. Soc. Am., 92 (6), 2369-2380.

Baisch, S., R. Weidler, R. Vörös, D. Wyborn, and L. de Graaf (2006), Induced seismicity during the stimulation of a geothermal HFR reservoir in the Cooper Basin, Australia, B. Seismol. Soc. Am., 96 (6), 2242-2256.

Baisch S., and R. Vörös (2009), AP3000 Report – Induced Seismicity, Company: Q-con GmbH, Ref KBA00.

Baisch, S., R. Vörös, R. Weidler, and D. Wyborn (2009). Investigation of fault mechanisms during geothermal reservoir stimulation experiments in the Cooper Basin, Australia, B. Seismol. Soc. Am., 99 (1), 148-158.

Baisch, S., R. Vörös, E. Rothert, H. Stang, R. Jung, and R. Schellschmidt (2010). A numerical model for fluid injection induced seismicity at Soultz-sous-Forêts, Int. J. Rock Mech. Min., 47, 405-413.

Baria, R., J. Baumgärtner, A. Gérard, R. Jung, and J. Garnish (1999), European HDR research programme at Soultz-sous-Forêts (France) 1987-1996, Geothermics, 28, 655-669.

Baria, R., S. Michelet, J. Baumgärtner, B. Dyer, J. Nicholls, T. Hettkamp, et al. (2005), Creation and mapping of 5000 m deep HDR/HFR reservoir to produce electricity, Proceedings World Geothermal Congress 2005, Antalya, Turkey, 24-29 April 2005.

Barton, C. A., M. D. Zoback, and K. L. Burns (1988), In-situ stress orientation and magnitude at the Fenton geothermal site, New Mexico, determined from wellbore breakouts, Geoph. Res. Lett., 15 (5), 467-470.

Baumgärtner, J., R. Jung, A. Gérard, R. Baria, and J. Garnish (1996), The European HDR project at Soultz sous Forêts: Stimulation of the second deep well and first circulation experiments, Proceedings 21st Workshop on Geothermal Reservoir Engineering, Stanford University, Stanford, California, January 22-24 1996, SGP-TR-151.

Bertani, R. (2005), World geothermal power generation in the period 2001-2005, Geothermics, 34, 651-690.

Bianco, F., E. D. Pezzo, G. Saccorotti, and G. Ventura (2004), The role of hydrothermal fluids in triggering the July-August 2000 seismic swarm at Campi Flegrei, Italy: evidence from seismological and mesostructural data, J. Volcanol. Geoth. Res., 133, 229-246.

Bodvarsson, G. S., S. M. Benson, O. Sigurdsson, V. Stefansson, and E. T. Eliasson (1984a), The Krafla geothermal field, Iceland: 1. Analysis of well test data, Water Resour. Res., 20 (11), 1515-1530.





46

Bodvarsson, G.S., K. Pruess, V. Stefansson, and E. T. Eliasson (1984b), The Krafla geothermal field, Iceland: 3. The generating capacity of the field, Water Resour. Res., 20 (11), 1545-1559.

Bohnhoff, M., S. Baisch, and H.-P. Harjes (2004), Fault mechanisms of induced seismicity at the superdeep German Continental Deep Drilling Program (KTB) borehole and their relation to fault structure and stress field, J. Geophys. Res., 109, B02309, doi:10.1029/2003JB002528.

Bommer, J. J., S. Oates, J. M. Cepeda, C. Lindholm, J. Bird, R. Torres, G. Marroquín, and J. Rivas (2006), Control of hazard due to seismicity induced by a hot fractured rock geothermal project, Eng. Geol., 83, 287-306.

Bouchot, V., B. Sanjuan, H. Traineau, L. Guillou-Frottier, L. Thinon, J. M. Baltassat, H. Fabriol, B. Bourgeois, and E. Lasne (2010), Assessment of the Bouillante geothermal field (Guadeloupe, French West Indies): Toward a conceptual model of the high temperature geothermal system, Proceedings World Geothermal Congress 2010, Bali, Indonesia, 25-29 April 2010, 8 pp.

Brown, D. W., and D. V. Duchane (1999), Scientific progress on the Fenton Hill HDR project since 1983, Geothermics, 28, 591-601.

Brudy, M., M. D. Zoback, K. Fuchs, F. Rummel, and J. Baumgärtner (1997), Estimation of the complete stress tensor to 8 km depth in the KTB scientific drill holes: Implications for crustal strength, J. Geophys. Res., 102 (B8), 18,453-18,475.

Bruel, D. (1995), Heat extraction modelling from forced fluid flow through stimulated fractured rock masses: Application to the Rosemanowes Hot Dry Rock reservoir, Geothermics, 24 (3), 361-374.

Cavarretta, G., G. Gianelli, G. Scandiffio, and F. Tecce (1985), Evolution of the Latera geothermal system II: Metamorphic, hydrothermal mineral assemblages and fluid chemistry, J. Volcanol. Geoth. Res., 26, 337-364.

Cladouhos, T., S. Petty, G. Foulger, B. Julian, and M. Fehler (2010), Injection induced seismicity and geothermal energy, Geoth. Res. T, 34, 1213-1220.

D’Amore, F., and J. T. Mejia (1999), Chemical and physical reservoir parameters at initial conditions in Berlin geothermal field, El Salvador: A first assessment, Geothermics, 28, 45-73.

Deichmann, N., and K. F. Evans (2010), Injection-induced seismicity: Placing the problem in perspective, Keynote Talk, European Seismological Commission 2010, Montpellier, France, 6-10 September 2010.

De Natale, G., and C. Troise (2011), The ‘Campi Flegrei Deep Drilling Project’: from risk mitigation to renewable energy production, European Review, 19 (3), 337-353.

Denlinger, R. P., and C. G. Bufe (1982), Reservoir conditions related to induced seismicity at the Geysers steam reservoir, Northern California, B. Seismol. Soc. Am., 72 (4), 1317-1327.

De Vivo, B., H. E. Belkin, M. Barbieri, W. Chelini, P. Lattanzi, A. Lima, and L. Tolomeo (1989), The Campi Flegrei (Italy) geothermal system: A fluid inclusion study of the Mofete and San Vito fields, J. Volcanol. Geoth. Res., 36, 303-326.

Dezayes, C., A. Genter, and B. Valley (2010), Structure of the low permeable naturally fractured geothermal reservoir at Soultz, C. R. Geosci., 342, 517-530.

Dinske, C. (2010), Interpretation of fluid-induced seismicity at geothermal and hydrocarbon reservoirs of Basel and Cotton Valley, PhD thesis, Freie Universität Berlin, 151 pp.

Dorbath, L., N. Cuenot, A. Genter, and M. Frogneux (2009), Seismic response of the fractured and faulted granite of Soultz-sous-Forêts (France) to 5 km deep massive water injections, Geophys. J. Int., 177, 653-675.

Dyer, B.C., U. Schanz, T. Spillmann, F. Ladner, and M. O. Häring (2008), Microseismic imaging of a geothermal reservoir stimulation, TLE, 27, 856-869.

Elders, W. A., and G. O. Friðleifsson (2010), The science program of the Iceland Deep Drilling Project (IDDP): a study of supercritical geothermal resources, Proceedings World Geothermal Congress 2010, Bali, Indonesia, 25-29 April 2010, 9 pp.





47

Emmermann, R., and J. Lauterjung (1997), The German Continental Deep Drilling Program KTB: Overview and major results, J. Geophys. Res., 102 (B8), 18,179-18,201.

Evans, K. F., H. Moriya, H. Niitsuma, R. H. Jones, W. S. Phillips, A. Genter, J. Sausse, R. Jung, and R. Baria (2005), Microseismicity and permeability enhancement of hydrogeologic structures during massive fluid injections into granite at 3 km depth at the Soultz HDR site, Geophys. J. Int., 160, 388-412.

Evans, K. F., A. Zappone, T. Kraft, N. Deichmann, and F. Moia, (in press), A survey of the induced seismic responses to fluid injection in geothermal and CO2 reservoirs in Europe, Geothermics.

Fehler, M. C. (1989), Stress control of seismicity patterns observed during hydraulic fracturing experiments at the Fenton Hill Hot Dry Rock geothermal energy site, New Mexico, Int. J. Rock Mech. Min., 26, 211-219.

Fischer, T., and A. Guest (2011), Shear and tensile earthquakes caused by fluid injection, Geoph. Res. Lett., 38, L05307, doi:10.1029/2010GL045447.

Franzson, H., E. Gunnlaugsson, K. Árnason, K. Sæmundsson, B. Steingrímsson, and B. S. Harðarson (2010), The Hengill geothermal system, conceptual model and thermal evolution, Proceedings World Geothermal Congress 2010, Bali, Indonesia, 25-29 April 2010, 9 pp.

Friðleifsson, G. O., W. A. Elders (2005), The Iceland Deep Drilling Project: a search for deep unconventional geothermal resources, Geothermics, 34, 269-285.

Friðleifsson, G. O., A. Albertsson, and W. A. Elders (2010), Iceland Deep Drilling Project (IDDP) – 10 years later – Still and opportunity for international collaboration, Proceedings World geothermal Congress 2010, Bali, Indonesia, 25-29 April 2010, 5 pp.

Gale, J. F. W., R. M. Reed, and J. Holder (2007), Natural fractures in the Barnett Shale and their importance for hydraulic fracture treatments, AAPG Bull., 91 (4), 603-622.

Genter, A., K. Evans, N. Cuenot, D. Fritsch, and B. Sanjuan (2010), Contribution of the exploration of deep crystalline fractured reservoir of Soultz to the knowledge of enhanced geothermal systems (EGS), C. R. Geosci., 342, 502-516.

Grawinkel, A., and B. Stöckhert (1997), Hydrostatic pore fluid pressure to 9 km depth- Fluid inclusion evidence from the KTB deep drill hole, Geophys. Res. Lett., 24, 3,273-3,276.

Green, A. S. P., and R. Baria, (1987), Active seismics to determine reservoir characteristics of a hot dry rock geothermal system, 12th Workshop on Geothermal Reservoir Engineering, Stanford University, California, January 20-22 1987, 131-137.

Haimson, B. C., and C. Chang (2002), True triaxial strength of the KTB amphibolites under borehole wall conditions and its use to estimate the maximum horizontal stress, J. Geophys. Res., 107 (B10), 2257, doi:10.1029/20001JB000647.

Hardarson, B. S., G. M. Einarsson, B. R. Kristjánsson, G. Gunnarsson, H. M. Helgadóttir, H. Franzson, K. Árnason, K., Agústsson, and E. Gunnlaugsson (2010), Geothermal reinjection at the Hengill triple junction, SW Iceland, Proceedings World Geothermal Congress 2010, Bali, Indonesia, 25-29 April 2010, 7 pp.

Häring, M. O., F. Ladner, U. Schanz, and T. Spillmann (2007), Deep heat mining Basel, Preliminary results, Proceedings European Geothermal Congress 2007, Unterhachung, Germany, 30 May - 1 June 2007, 9 pp.

Häring, M. O., U. Schanz, F. Ladner, and B. C. Dyer (2008), Characterisation of the Basel 1 enhanced geothermal system, Geothermics, 37, 469-495.

Huenges, E., J. Erzinger, J. Kück, B. Engeser, and W. Kessels (1997), The permeable crust: Geohydraulic properties down to 9101 m depth, J. Geophys. Res., 102 (B8), 18,255-18,265.

IEA (2003), Geothermal annual report. Ito, T., R. Fuji, K. F. Evans, and K. Hayashi, 2002, Estimation of stress profile with depth

from analysis of temperature and fracture orientation logs in a 3.6 km deep well at Soultz, France, SPE/ISRM Rock Mechanics Conference, Irving, Texas, 20-23 October 2002.

Kadko, D., K. Gronvold, and D. Butterfield (2007), Application of radium isotopes to determine crustal residence times of hydrothermal fluids from two sites on the Reykjanes Peninsula, Iceland, Geochim. Cosmochim. Ac., 71, 6019-6029.





48

Kahn, D.S. (2008), Hydro-fractured reservoirs: A study using double-difference location techniques, PhD Thesis, Duke University, 175 pp.

Kaieda, H., S. Sasaki, and D. Wyborn (2010), Comparison of characteristics of micro-earthquakes observed during hydraulic stimulation operations in Ogachi, Hijiori and Cooper Basin HDR projects, Proceedings World Geothermal Congress 2010, Bali, Indonesia, 25-29 April 2010, 6 pp.

Karner, S. L. (2005), Stimulation techniques used in Enhanced Geothermal Systems: Perspectives from geomechanics and rock physics, Proceedings 30th Workshop on Geothermal Reservoir Engineering, Stanford University, Stanford, California, US, January 31 - February 2 2005, SGP-TR-176, 9 pp.

Küper, L. (2004), Source parameters from fluid-injection-induced microearthquakes at the KTB, Master thesis, Ruhr-University Bochum.

Kwiatek, G., M. Bohnhoff, G. Dresen, A. Schulze, T. Schulte, G. Zimmermann, and E. Huenges (2008), Microseismic event analysis in conjunction with stimulation treatments at the geothermal research well GtGrSk4/05 in Groβ Schönebeck/Germany, Proceedings of the 33rd Workshop on Geothermal Reservoir Engineering, Stanford University, Stanford, CA, US, January 28-30, SGP-TR-185, 7 pp.

Kwiatek, G., M. Bohnhoff, G. Dresen, A. Schulze, T. Schulte, G. Zimmermann, and E. Huenges (2010), Microseismicity induced during fluid-injection: A case study from the geothermal site at Groβ Schönebeck, North German Basin, Acta Geophys, 58 (6), 995-1020.

Laughlin, A. W., A. C. Eddy, R. Laney, and M. J. Aldrich (1983), Geology of the Fenton Hill, New Mexico, Hot Dry Rock site, J. Volcanol. Geoth. Res., 15, 21-41.

Lipman, S. C., C. J. Strobel, and M. S. Gulati (1978), Reservoir performance of The Geysers Field, Geothermics, 7, 209-219.

Lopez, S., V. Bouchot, M. Lakhssassi, P. Calcagno, and B. Grappe (2010), Modeling of Bouillante geothermal field (Guadeloupe, French Lesser Antilles), Proceedings 35th Workshop on Geothermal Reservoir Engineering, Stanford University, Stanford, California, February 1-3 2010, 8 pp.

Mahrer, K., J. Ake, L. Block, D. O’Connell, and J. Bundy (2005), Injecting brine and inducing seismicity at the world’s deepest injection well, Paradox Valley, Southwest Colorado, Dev. Water Sci., 52, 361-375.

Majer, E. L., R. Baria, M. Stark, S. Oates, J. Bommer, B. Smith, and H. Asanuma (2007), Induced seismicity associated with Enhanced Geothermal Systems, Geothermics, 36, 185-222.

Majer, E. L., T. V. McEvilly (1979), Seismological investigations at The Geysers Geothermal Field, Geophysics, 44, 246-269.

Majer, E. L., and J. E. Peterson (2007), The impact of injection on seismicity at The Geysers, California Geothermal Field, Int. J. Rock Mech. Min., 44, 1079-1090.

Mas, A., D. Guisseau, P. P. Mas, D. Beaufort, A. Genter, B. Sanjuan, and J. P. Girard (2006), Clay minerals related to the hydrothermal activity of the Bouillante geothermal field (Guadeloupe), J. Volcanol. Geoth. Res., 158, 380-400.

Moeck, I., G. Kwiatek, and G. Zimmermann (2009), Slip tendency analysis, fault reactivation potential and induced seismicity in a deep geothermal reservoir, J. Struct. Geol., 31, 1174-1182.

Moriya, H., K. Nakazato, H. Niitsuma, and R. Baria (2002), Detailed fracture system of the Soultz-sous-Forêts HDR field evaluated using microseismic multiplet analysis, Pure Appl. Geophys., 159, 517-541.

Mossop, A., and P. Segall (unpublished study), Induced seismicity in geothermal fields II – Correlation and interpretation at The Geysers.

Nathenson, M. (1999), The dependence of permeability on effective stress from flow tests at hot dry rock reservoirs at Rosemanowes (Cornwall) and Fenton Hill (New Mexico), Geothermics, 28, 315-340.





49

Oates, S., B. Hustedt, G. Siddiqi, B. Worrall, and M. Darnet (2004), Microseismic monitoring of the HFR hydraulic stimulation of well TR8A in the Berlin geothermal field, El Salvador, Shell Internal Report, 91 pp.

O’Connell, D. R. H., and L. R. Johnson (1991), Progressive inversion for hypocenters and P wave and S wave velocity structure: Application to the Geysers, California, geothermal field, J. Geophys. Res., 96 (B4), 6,223-6,236.

Oppenheimer, D. H. (1986), Extensional tectonics at The Geysers geothermal area, California, J. Geophy. Res., 91 (B11), 11,463-11,476.

Ostermann, D. (2008), Earthquake risks resulting from geothermal power plant-reasons for the earthquake series in Basel in 2007, JSEE, 10 (1), 49-52.

Parker, R.H. (1991), Problems in the development of artificial geothermal energy exploitation in Cornwall, Proceedings of the Ussher Society, 7, 316-320.

Parker, R. (1999), The Rosemanowes HDR project 1983-1991, Geothermics, 28, 603-615.

Pauwels, H. (1997), Geochemical results of a single-well hydraulic injection test in an experimental Hot Dry Rock geothermal reservoir, Soultz-sous-Forêts, Alsace, France, Appl. Geochem., 12, 661-673.

Phillips, W. S., J. T. Rutledge, L. S. House, and M. C. Fehler (2002), Induced microearthquake patterns in hydrocarbon and geothermal reservoirs: Six case studies, Pure Appl. Geophys., 159, 345-369.

Pine, R. J., P. Ledingham, and C. M. Merrifield (1983), In-situ stress measurements in the Carnmenellis granite - II. Hydrofracture tests at Rosemanowes Quarry to depths of 2000 m, Int. J. Rock Mech. Min., 20 (2), 63-72.

Raymond, J., A. E. Williams-Jones, and J. R. Clark (2005), Mineralization associated with scale and altered rock and pipe fragments from the Berlín geothermal field, El Salvador; implications for metal transport in natural systems, J. Volcanol. Geoth. Res., 145, 81-96.

Regenspurg, S., T. Wiersberg, W. Brandt, E. Huenges, A. Saadat, K. Schmidt, and G. Zimmermann (2010), Geochemical properties of saline geothermal fluids from the in-situ geothermal laboratory Groβ Schönebeck (Germany), Chem. Erde-Geochem., 70 (S3), 3-12.

Reynolds, S. D., S. D. Mildren, R. R. Hillis, J. J. Meyer, and T. Flottmann (2005), Maximum horizontal stress orientations in the Cooper Basin, Australia: implications for plate-scale tectonics and local stress sources, Geophys. J. Int., 160, 331-343.

Reynolds, S. D., S. D. Mildren, R. R. Hillis, J. J. Meyer (2006), Constraining stress magnitudes using petroleum exploration data in the Cooper-Eromanga Basins, Australia, Tectonophysics, 415, 123-140.

Richards, H. G. (1992), Granite-water reactions in an experimental Hot Dry Rock geothermal reservoir, Rosemanowes test site, Cornwall, U.K., Appl. Geochem., 7, 193-222.

Richards, H. G., R. H. Parker, A. S. P. Green, R. H. Jones, J. D. M. Nicholls, D. A. C. Nicol, et al. (1994), The performance and characteristics of the experimental hot dry rock geothermal reservoir at Rosemanowes, Cornwall (1985-1988), Geothermics, 23 (2), 73-109.

Rodríguez, J. A. (2000), Geothermal development in El Salvador – A country update, Proceedings World Geothermal Congress 2000, Kyushu – Tohoku, Japan, May 28 - June 10 2000.’

Ruggieri, G., C. M. Petrone, G. Gianelli, A. Arias, and E. T. Henriquez (2006), Hydrothermal alteration in the Berlin geothermal field (El Salvador): new data and discussion on the natural state of the system, Period. Mineral., 75, 293-312.

Rutledge, J. T., and R. J. Zinno (1998), Microseismic mapping of a Cotton Valley hydraulic fracture using decimated downhole arrays, Proceedings 68th Annual International Meeting, SEG Expanded Abstracts, 338-341.

Rutledge, J. T., and W. S. Phillips (2003), Hydraulic stimulation of natural fractures as revealed by induced microearthquakes, Carthage Cotton Valley gas field, East Texas, Geophysics, 68 (2), 441-452.





50

Rutledge, J.T., W. S. Phillips, and M. J. Mayerhofer (2004), Faulting induced by forced fluid injection and fluid flow forced by faulting: An interpretation of hydraulic-fracture microseismicity, Carthage Cotton Valley gas field, Texas, B. Seismol. Soc. Am., 94 (5), 1817-1830.

Rutqvist, J., and C. Oldenburg (2007), Analysis of cause and mechanism for injection-indued seismicity at the Geysers Geothermal Field, California, Geoth. Res. T., 31, 441-445.

Sanjuan, B., P. Jousset, G. Pajot, N. Debeglia, M. De Michele, M. Brach, F. Dupont, G. Braibant, E. Lasne, and F. Duré (2010a), Monitoring of the Bouillante geothermal exploitation (Guadeloupe, French West Indies) and the impact on its immediate environment, Proceedings World Geothermal Congress 2010, Bali, Indonesia, 25-29 April 2010, 11 pp.

Sanjuan, B., R. Millot, M. Brach, R. Asmundsson, and N. Giroud (2010b), Use of a new sodium/lithium (Na/Li) geothermometric relationship for high-temperature dilute geothermal fluids from Iceland, Proceedings World Geothermal Congress 2010, Bali, Indonesia, 25-29 April 2010, 12 pp.

Shapiro, S. A., C. Dinske, and J. Kummerow (2007), Probability of a given-magnitude earthquake induced by a fluid injection, Geophys. Res. Lett., 34, L22314, doi:10.1029/2007GL031615.

Shapiro, S. A. (2008), Microseismicity: A tool for reservoir characterization, Eur. Assoc. of Geosci. and Eng., Houten, Netherlands.

Soma, N., H. Asanuma, H. Kaieda, K. Tezuka, D. Wyborn, and H. Niitsuma (2004), On site mapping of microseismicity at Cooper Basin, Australia HDR project by the Japanese team, Proceedings 29th Workshop on Geothermal Reservoir Engineering, Stanford University, Stanford, California, January 26-28 2004, SGP-TR-175.

Sveinbjornsdottir, A. E., M. L. Coleman, and B. W. D. Yardley (1986), Origin and history of hydrothermal fluids of the Reykjanes and Krafla geothermal fields, Iceland: A stable isotope study, Contrib. Mineral Petr., 94, 99-109.

Tassi, F., O. Vaselli, B. Capaccioni, G. Montegrossi, F. Barahona, and A. Caprai (2007), Scrubbing process and chemical equilibria controlling the composition of light hydrocarbons in natural gas discharges: An example from the geothermal fields of El Salvador, Geochem. Geophy. Geosy., 8 (5), Q05008, doi:10.1029/2006GC001487.

Tester, J.W., B. J. Anderson, A. S. Batchelor, D. D. Blackwell, R. DiPippo, E. M. Drake et al. (2006), Chapter 4 Review of EGS and related technology –Status and achievements: The Future of Geothermal Energy - Impact of Enhanced Geothermal Systems (EGS) on the United States in the 21st century, Massachusetts Institute of Tecnhology, Cambridge, MA, USA, 372 pp.

Traineau, H., B. Sanjuan, D. Beaufort, M. Brach, C. Castaing, H. Correia, A. Genter, and B. Herbrich (1997), The Bouillante geothermal field (F.W. I.) revisited: New data on the fractured geothermal reservoir in light of a future stimulation experiment in a low productive well, Proceedings 22nd Workshop on Geothermal Reservoir Engineering, Stanford University, Stanford, California, January 27-29 1997, SGP-TR-155, 97-103.

Urpi, L., G. Zimmermann, G. Blöcher, and G. Kwiatek (2011), Microseismicity at Groβ Schönebeck – A case review, Proceedings of the 36th Workshop on Geothermal Reservoir Engineering, Stanford University, Stanford, California, January 31 - February 2 2011, SGP-TR-191, 8 pp.

Valentino, G. M., G. Cortecci, E. Franco, and D. Stanzione (1999), Chemical and isotopic compositions of minerals and waters from the Campi Flegrei volcanic system, Naples, Italy, J. Volcanol. Geoth. Res, 91, 329-344.

Valley, B. C. (2007), The relation between natural fracturing and stress heterogeneities in deep-seated crystalline rocks at Soultz-sous-Forêts (France), PhD thesis, ETH-Zürich, Switzerland, 260 pp.

Valley, B., and K. F. Evans (2009), Stress orientation to 5 km depth in the basement below Basel (Switzerland) from borehole failure analysis, Swiss. J. Geosci., 102, 467-480.





51

Van Eijs, R. M. H. E., F. M. M. Mulders, M. Nepveu, C. J. Kenter, and B. C. Scheffers (2006), Correlation between hydrocarbon reservoir properties and induced seismicity in the Netherlands, Eng. Geol., 84, 99-111.

Zang, A., and O. Stephansson (2010), Stress field of the Earth’s crust, doi:10.1007/987-1-4020-8444-7_9.

Zimmermann, G., I. Moeck, and G. Blöcher (2010), Cyclic waterfrac stimulation to develop an enhanced geothermal system (EGS): Conceptual design and experimental results, Geothermics, 39, 59-69.





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4.6 Appendix I

This appendix gives part of the data which has been collected in this study (see Excel sheet for complete overview of all the parameters).The tables below include parameters related to the depth of injection and bottom hole temperature (Table 1), the type of reservoir rock (Table 2), the induced seismicity (Table 3), the injections (Table 4), stress and pressure in reservoir (Table 5) and fracture network (Table 6).





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Table 1 Type of site, true vertical depth (TVD) of injection well and bottom hole temperature (BHT) for the selected sites. NF = not found. Sites; injection year Type of site TVD [m] BHT [°C] Soultz-sous forêts; GPK-1 (1993) EGS 3590 160 Soultz-sous forêts; GPK-2 (1995) EGS 3876 168 Soultz-sous forêts; GPK-2 (1996) EGS 3876 168 Soultz-sous forêts; GPK-2 (2000) EGS 4955 202 Soultz-sous forêts; GPK-3 (2003) EGS 5091 200.6 Soultz-sous forêts; GPK-4 (2004 & 2005) EGS 4982 200 Basel; Basel-1 (2006) EGS 5009 200 Krafla (Iceland) EGS 1000-2400 190-210; 240-340

Reykjanes (Iceland) EGS 1000- > 2500 260-310

Hengill (Iceland) EGS 1000-3322 200-320 Berlín; TR8-A (2002-2003) EGS 2465.52 305 Latera; Latera-1 (1981 & 1982) EGS ~2783 ~350 Latera; Latera-6 (1981) EGS ~2017 ~225 The Geysers EGS NF ~240 Groβ Schönebeck; sand (2007) EGS ~4285 150 Groβ Schönebeck; volcanic (2007) EGS ~4285 150 KTB; KTB-HB (1994) EGS 9100 265 KTB; KTB-HB (2000) EGS 9100 265 Rosemanowes; RH12 (1982) EGS 2115 79 Rosemanowes; RH15 (1985) EGS 2650 99.8 Rosemanowes; RH15 (1990) EGS 2650 99.8 Cooper Basin; Habanero-1 (2003) EGS 4421 250 Cooper Basin; Habanero-1 (2005) EGS 4421 250 Bouillante EGS 800 250-260 Fenton Hill; EE-2 (1983) EGS 4390 323 Fenton Hill; EE-3 (1984) EGS 3980 NF Fenton Hill; EE-3a (1985) EGS 4018 265 Fenton Hill; EE-3a (1986) EGS 4018 265 Paradox Valley; (1991-1995) Brine injection 4900 NF Paradox Valley; (1996-1999) Brine injection 4900 NF Paradox Valley; (1999-2000) Brine injection 4900 NF Paradox Valley; (2000-2002) Brine injection 4900 NF Paradox Valley; (2002--2003) Brine injection 4900 NF Ogachi; OGC-1 (1991) EGS 1000 230 Cotton Valley; well 21-10 (1997) Hydraulic fracture ~2980 120 Cotton Valley; well 21-10 (1997) Hydraulic fracture ~2980 120 Cotton Valley; well 21-09 (1997) Hydraulic fracture ~2780 120 Cotton Valley; well 21-09 (1997) Hydraulic fracture ~2780 120 Cotton Valley; well 21-09 (1997) Hydraulic fracture ~2780 120 Barnett shale Hydraulic fracture NF NF Campi Flegrei EGS NF 350-420





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Table 2 Rock type and static Young’s modulus (E), static Poisson’s ratio (ν), porosity (φ) and permeability (k) for reservoir rock. NF= not found.

Sites; injection year Rock type E [GPa] ν φ [%] k

Soultz-sous forêts; GPK-1 (1993) Granite 54 ± 2 0.2-0.25 NF NF Soultz-sous forêts; GPK-2 (1995) Granite 54 ± 2 0.2-0.25 NF NF Soultz-sous forêts; GPK-2 (1996) Granite 54 ± 2 0.2-0.25 NF NF Soultz-sous forêts; GPK-2 (2000) Granite 54 ± 2 0.2-0.25 NF NF Soultz-sous forêts; GPK-3 (2003) Granite 54 ± 2 0.2-0.25 NF NF Soultz-sous forêts; GPK-4 (2004 & 2005) Granite 54 ± 2 0.2-0.25 NF NF

Basel; Basel-1 (2006) Granite NF NF 1 ~10 µD; ~150 µD

Krafla (Iceland) Volcanic NF NF 2-5 NF Reykjanes (Iceland) Volcanic NF NF NF NF Hengill (Iceland) Volcanic NF NF NF NF Berlín; TR8-A (2002-2003) Volcanic NF NF NF ~100 mD

Latera; Latera-1 (1981 & 1982) Limestones, igneous rock NF NF NF NF

Latera; Latera-6 (1981) Limestones, igneous rock NF NF NF NF

The Geysers Graywacke 33 0.24 NF 1-100 µD Groβ Schönebeck; sand (2007) Sandstone ~55 0.18 8-10 16.5 mD Groβ Schönebeck; volcanic (2007) Volcanic ~55 0.2 0.10 Low

KTB; KTB-HB (1994) Gneisses, amphibolites 95 ± 5

0.26 ± 0.03

0.70 ± 0.33 ~8 µD

KTB; KTB-HB (2000) Gneisses, amphibolites 95 ± 5

0.26 ± 0.03

0.70 ± 0.33 ~8 µD

Rosemanowes; RH12 (1982) Granite 61.1 0.22-0.27 NF 10-100 µD Rosemanowes; RH15 (1985) Granite 62.7 0.22-0.27 NF 10-100 µD Rosemanowes; RH15 (1990) Granite 62.7 0.22-0.27 NF 10-100 µD Cooper Basin; Habanero-1 (2003) Granite NF NF NF NF Cooper Basin; Habanero-1 (2005) Granite NF NF NF NF Bouillante Volcanic NF NF NF NF Fenton Hill; EE-2 (1983) Granite NF NF NF low Fenton Hill; EE-3 (1984) Granite NF NF NF low Fenton Hill; EE-3a (1985) Granite NF NF NF low Fenton Hill; EE-3a (1986) Granite NF NF NF low

Paradox Valley; (1991-1995) Dolomitic limestone NF NF < 6 NF




Paradox Valley; (2002--2003) Dolomitic limestone NF NF < 6 NF

Ogachi; OGC-1 (1991) Granodiorite NF NF NF 0.1-1 mD; 10-100 mD





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Cotton Valley; well 21-10 (1997) Gas-bearing sands NF NF

≤ 10; 30 ~1 µD

Cotton Valley; well 21-10 (1997) Gas-bearing sands NF NF

≤ 10; 30 ~1 µD

Cotton Valley; well 21-09 (1997) Gas-bearing sands NF NF ≤ 10 ~1 µD



Barnett shale Gas-bearing shale rock 33 0.2-0.3 3-5 0.1-0.5 µD

Campi Flegrei Volcanic NF NF NF NF





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Table 3 Seismicity parameters of the studied sites. Maximum observed magnitude (Mmax) is the local magnitude. It is not certain during which experiment at Rosemanowes and Fenton Hill the Mmax listed in this table (values in pink) has been induced. NF= not found.

Sites; injection year Range of magnitudes # events Mmax Delay

Soultz-sous forêts; GPK-1 (1993) -2 >Mw < 1 ~12000 2 ~17 hr Soultz-sous forêts; GPK-2 (1995) NF NF NF NF Soultz-sous forêts; GPK-2 (1996) NF NF NF NF Soultz-sous forêts; GPK-2 (2000) NF > 30000 2.5 no delay Soultz-sous forêts; GPK-3 (2003) NF 90648 2.9 no delay

Soultz-sous forêts; GPK-4 (2004 & 2005) NF 22718

2.3 (2004); 2.7 (2005) no delay

Basel; Basel-1 (2006) NF > 10500 3.4 ~3.75 hr Krafla (Iceland) NF NF NF NF Reykjanes (Iceland) NF NF NF NF Hengill (Iceland) NF NF NF NF Berlín; TR8-A (2002-2003) NF 32 (1), 49 (2), 24 (3) 4.4 NF

Latera; Latera-1 (1981 & 1982) NF 193 (1);147 (2); 372 (3)

0.4 (1); 0.2 (2); 0.5 (3)

~3hr (1); ~56 hr (2); ~103 hr (3)

Latera; Latera-6 (1981) NF 20 (1); 90 (2);191 (3)

-0.4 (1); -0.1 (2); 0.4 (3)

~5 hr (1); ~26 hr (2); ~11 hr (3)

The Geysers NF NF 4.6 (1982) NF

Groβ Schönebeck; sand (2007) NF none NF NF

Groβ Schönebeck; volcanic (2007) -1.9 < Mw < -1.1 70 -1.1 (Mw) ~20 min

KTB; KTB-HB (1994) NF 400 1.2 NF KTB; KTB-HB (2000) NF 2799 0.7 NF Rosemanowes; RH12 (1982) NF NF 2 NF Rosemanowes; RH15 (1985) NF 270 2 NF Rosemanowes; RH15 (1990) NF NF 2 NF Cooper Basin; Habanero-1 (2003) NF 27000 3.7 no delay Cooper Basin; Habanero-1 (2005) -1.2 < Ml < 2.9 16000 2.9 22 hr Bouillante - none - - Fenton Hill; EE-2 (1983) NF 11366 1.5 NF Fenton Hill; EE-3 (1984) NF 946 1.5 NF Fenton Hill; EE-3a (1985) NF 1021 1.5 NF Fenton Hill; EE-3a (1986) NF 1962 1.5 NF Paradox Valley; (1991-1995) NF 666 NF NF Paradox Valley; (1996-1999) NF 2446 3.6 NF Paradox Valley; (1999-2000) NF 496 4.3 NF Paradox Valley; (2000-2002) NF 140 2.8 NF Paradox Valley; (2002--2003) NF 277 NF NF Ogachi; OGC-1 (1991) NF 1553 2 NF Cotton Valley; well 21-10 (1997) NF 1122 NF NF Cotton Valley; well 21-10 (1997) NF ~1200 NF NF





57

Cotton Valley; well 21-09 (1997) NF NF NF NF Cotton Valley; well 21-09 (1997) NF NF NF NF Cotton Valley; well 21-09 (1997) NF NF NF NF Barnett shale NF ~900 NF NF Campi Flegrei NF NF NF NF Table 4 Duration of injection (Tinj), maximum wellhead pressure (Pwh), injected volume (Vinj) and maximum injection rate (Qmax). NF= not found.

Sites; injection year Tinj Pwh [MPa] Vinj [m3] Qmax [L s-1]

Soultz-sous forêts; GPK-1 (1993) NF NF 25000 36 Soultz-sous forêts; GPK-2 (1995) NF 12.3 28000 56 Soultz-sous forêts; GPK-2 (1996) NF 13 28000 78 Soultz-sous forêts; GPK-2 (2000) 141 hr 14.5 22680 50 Soultz-sous forêts; GPK-3 (2003) 255 hr 16 37300 100

Soultz-sous forêts; GPK-4 (2004 & 2005)

83 hr (2004); 95 hr (2005)

17 (2004); 14 (2005)

9300 (2004); 12300 (2005)

45 (2004); 45 (2005)

Basel; Basel-1 (2006) 6 d ~30 11500 60 Krafla (Iceland) NF NF NF NF Reykjanes (Iceland) NF NF NF NF Hengill (Iceland) NF NF NF NF

Berlín; TR8-A (2002-2003) NF

~7.8 (1); ~9 (2); ~13.5 (3) 300000

~108 (1); ~165 (2); ~120 (3)

Latera; Latera-1 (1981 & 1982)

~53 hr (1); ~100 hr (2); ~43 hr (3)

7.2 (1); 6.7 (2); 9.5 (3) NF

16.7 (1); 25.6 (2); 125 (3)

Latera; Latera-6 (1981)

~16 hr (1); ~40 hr (2); ~19 hr (3)

7.6 (1); 14.6 (2); 14.2 (3) NF

66.7 (1); 27.8 (2); 40.3 (3)

The Geysers NF NF NF 1997: 255 2003: 347

Groβ Schönebeck, sand (2007) NF 49.5; 38 500 gel 66; 58 Groβ Schönebeck, volcanic (2007) 6 d 58.6 13000 150

KTB; KTB-HB (1994) 24 hr 50 210 CaBr2/KCl2 9.2

KTB; KTB-HB (2000) 60 d 30 4000 1.2 Rosemanowes; RH12 (1982) NF 14 30000 100

Rosemanowes; RH15 (1985) 8 hr (gel)

15 (gel); 16.3 (water)

5500 gel; 200 water 260

Rosemanowes; RH15 (1990) NF NF 4000 (low viscosity gel) NF

Cooper Basin; Habanero-1 (2003) NF 65.5 20000 48 Cooper Basin; Habanero-1 (2005) 13 d 62 22500 > 31

Bouillante - -

4.5 million tons produced per year

600 tons/h produced water

Fenton Hill; EE-2 (1983) ~62 hr ~48 21600 100





58

Fenton Hill; EE-3 (1984) NF NF 7600 20 Fenton Hill; EE-3a (1985) NF NF 5600 20 Fenton Hill; EE-3a (1986) NF NF 4400 20 Paradox Valley; (1991-1995) 438 d 32 620072 25 Paradox Valley; (1996-1999) ~1100 d 33.8 NF 21.5 Paradox Valley; (1999-2000) ~332 d 33.8 NF 21.5 Paradox Valley; (2000-2002) ~566 d 30.3 NF 14.3 Paradox Valley; (2002--2003) ~724 d 30.3 NF 14.3 Ogachi; OGC-1 (1991) ~11 d ~19 > 10000 12.5 Cotton Valley; well 21-10 (1997) NF NF 1340 119 Cotton Valley; well 21-10 (1997) NF NF 1253 106 Cotton Valley; well 21-09 (1997) NF NF 419 26.5 Cotton Valley; well 21-09 (1997) NF NF 396 26.5 Cotton Valley; well 21-09 (1997) NF NF 400 26.5 Barnett shale ~5.5 hr 42 2840 150 Campi Flegrei NF NF NF NF





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Table 5 Magnitude of vertical stress (Sv), magnitude and orientation of minimum horizontal stress (Sh) and maximum horizontal stress (SH) and maximum formation pressure (Pformation). NF= not found.

Sites; injection year Sv [MPa]

Sh [MPa]

Orientation Sh

SH [MPa]

Orientation SH

Pformation [MPa]

Soultz-sous forêts; GPK-1 (1993) 71.4 36.5 NF 70.5 N170° E ~38 Soultz-sous forêts; GPK-2 (1995) 80.6 41.9 NF 82.6 NF 45 Soultz-sous forêts; GPK-2 (1996) 80.6 41.9 NF 82.6 NF NF

Soultz-sous forêts; GPK-2 (2000) 111.0 ± 4.9

60.2 ± 3.8 260°

111 +22.2/ -33.3

N169°E ± 14° ~56.9

Soultz-sous forêts; GPK-3 (2003) 113.1 ± 5.0

61.5 ± 3.8 260°

113.1 +22.6/ -33.9

N169°E ± 14° 61.5

Soultz-sous forêts; GPK-4 (2004 & 2005)

113.2 ± 5.0

61.5 ± 3.8 260°

113.2 +22.6/ -34.0

N169°E ± 14° 62.1

Basel; Basel-1 (2006) 124.7 ~84.2 54° ± 14° ~163.2 - 257.9

N144° E ± 14° 74

Krafla (Iceland) NF NF NF NF NF NF Reykjanes (Iceland) NF NF NF NF NF NF Hengill (Iceland) NF NF NF NF NF NF Berlín; TR8-A (2002-2003) NF NF NF NF NF NF Latera; Latera-1 (1981 & 1982) NF NF NF NF NF NF Latera; Latera-6 (1981) NF NF NF NF NF NF The Geysers NF NF NF NF NF NF

Groβ Schönebeck, sand (2007) 100 ~55 108.5° ± 3.7° ~98

018.5° ± 3.7° ~40

Groβ Schönebeck, volcanic (2007) 103 ~72 108.5° ± 3.7° ~105

018.5° ± 3.7° 86

KTB; KTB-HB (1994) ~278.1 ~156.3 NF ~300 N160° ± 10° E NF

KTB; KTB-HB (2000) ~278.1 ~156.3 NF ~300 N160° ± 10° E 120

Rosemanowes; RH12 (1982) 46.1 27.3 NF 64.6 N130° NF Rosemanowes; RH15 (1985) 56.7 32.2 NF 76.1 N130° 35.3 Rosemanowes; RH15 (1990) 56.7 32.2 NF 76.1 N130° NF

Cooper Basin; Habanero-1 (2003) 86.3-97.2 64.1 NF

111.6-215.6 101° N ~110

Cooper Basin; Habanero-1 (2005) 86.3-97.2 64.1 NF

111.6-215.6 101° N NF

Bouillante NF NF NF NF NF NF Fenton Hill; EE-2 (1983) NF NF NF NF NF 48 Fenton Hill; EE-3 (1984) NF NF 119° NF NF NF Fenton Hill; EE-3a (1985) NF NF NF NF NF NF Fenton Hill; EE-3a (1986) NF NF NF NF NF NF Paradox Valley; (1991-1995) NF NF NF NF NF NF Paradox Valley; (1996-1999) NF NF NF NF NF 80.7 Paradox Valley; (1999-2000) NF NF NF NF NF 80.7 Paradox Valley; (2000-2002) NF NF NF NF NF 77.2





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Paradox Valley; (2002--2003) NF NF NF NF NF 79.3 Ogachi; OGC-1 (1991) NF NF NF NF NF NF Cotton Valley; well 21-10 (1997) 63 33 NNW NF 80° 43 Cotton Valley; well 21-10 (1997) 66 38 NNW NF 80° 47 Cotton Valley; well 21-09 (1997) 63 33 NNW NF 80° 52 Cotton Valley; well 21-09 (1997) 63 33 NNW NF 80° 52 Cotton Valley; well 21-09 (1997) 66 38 NNW NF 80° 54 Barnett shale NF NF NF NF NE-SW NF

Campi Flegrei NF NF Strike 45°; plunge 8° NF

Strike 316°; plunge 7° NF





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Table 6 Orientation, length, width, height and volume of seismic cloud (Vseism) (fracture network). The dimensions of the seismic cloud (fracture network) are only estimates. Values in green are estimated from figure 3 and 4 in Evans et al. (in press). NF= not found.

Sites; injection year Orientation Length [m]

Width [m]

Height [m]

Vseism [m3]

Soultz-sous forêts; GPK-1 (1993)

25° NW; nearly vertical 1200 500 1500 9*108

Soultz-sous forêts; GPK-2 (1995) NF NF NF NF 2.4*108 Soultz-sous forêts; GPK-2 (1996) NF NF NF NF NF

Soultz-sous forêts; GPK-2 (2000) N20° W; dip: W72° 1600 500 1200 9.6*108

Soultz-sous forêts; GPK-3 (2003)

N10° W (main cloud); N20° W (northern extension) ~2100 ~800 ~1400 2.352*109

Soultz-sous forêts; GPK-4 (2004 & 2005) NS to 10° E

~800 (2004); ~1000 (2005)

~500 (2004); ~800 (2005)

~700 (2004); ~1000 (2005)

2.8*108 (2004); 8*108 (2005)

Basel; Basel-1 (2006) N159° E, subvertical ~1000 ~200 ~1000 2*108

Krafla (Iceland) NF NF NF NF NF Reykjanes (Iceland) NF NF NF NF NF Hengill (Iceland) NF NF NF NF NF Berlín; TR8-A (2002-2003) NF ~3200 ~1200 ~1600 6.1*109 Latera;Latera-1 (1981 & 1982) NF ~1900 ~600 ~2200 2.508*109 Latera; Latera-6 (1981) NF ~1400 ~700 ~2600 2.548*109 The Geysers NF NF NF NF NF Groβ Schönebeck, sand (2007) NF NF NF NF NF

Groβ Schönebeck, volcanic (2007)

017° (±10°); dip: 52° SE (±5°) ~300 ~100 ~300 9*106

KTB; KTB-HB (1994) ~SW-NE ~1600 ~750 ~880 1.1*109

KTB; KTB-HB (2000)

~ SE-NW (5-6 km); ~NS (9 km)

~540 (5-6 km); ~650 (9 km)

~260 (5-6 km); ~400 (9 km)

~560 (5-6 km); ~350 (9 km)

7.9*107 (5-6 km); 9.1*107 (9 km)

Rosemanowes; RH12 (1982) NF NF NF NF NF Rosemanowes; RH15 (1985) vertical 70 70 200 9.8*105 Rosemanowes; RH15 (1990) NF NF NF NF NF Cooper Basin; Habanero-1 (2003) dip < 10° ~3000 ~1000 ~350 1.05*109 Cooper Basin; Habanero-1 (2005) Subhorizontal NF NF ~150 NF

Bouillante

N-S fracture network connecting E-W faults NF NF

2500-3000 NF

Fenton Hill; EE-2 (1983) N10° W, dip: 65° E 1000 300 1000 3*108





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Fenton Hill; EE-3 (1984) NF NF NF NF NF Fenton Hill; EE-3a (1985) NF NF NF NF NF Fenton Hill; EE-3a (1986) NF NF NF NF NF Paradox Valley; (1991-1995) dip:15° E ~5400 ~3500 ~2100 3.969*109 Paradox Valley; (1996-1999) dip:15° E ~5400 ~3500 ~2100 3.969*109 Paradox Valley; (1999-2000) dip:15° E ~5400 ~3500 ~2100 3.969*109 Paradox Valley; (2000-2002) dip:15° E ~5400 ~3500 ~2100 3.969*109 Paradox Valley; (2002--2003) dip:15° E ~5400 ~3500 ~2100 3.969*109 Ogachi; OGC-1 (1991) NNE ~1000 ~400 ~500 2*108 Cotton Valley; well 21-10 (1997) N80° E ~540 ~53 ~120 3.4*106

Cotton Valley; well 21-10 (1997) N70° E; dip: 45° SE; ~390 ~61 ~87 2.1*106

Cotton Valley; well 21-09 (1997) N80° E ~383 ~25 ~50 4.6*105 Cotton Valley; well 21-09 (1997) NF NF NF NF NF Cotton Valley; well 21-09 (1997) N80° E ~530 ~50 ~25 6.6*105 Barnett shale NF ~833 ~267 ~100 2.224*107 Campi Flegrei NF NF NF NF NF

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4.7 Appendix II Definition of parameters listed in Appendix I/Excel sheet)

• Type of site: For each site it is mentioned what kind of site it is; most of the sites are

geothermal. The type of treatment is given between brackets. • Depth: This is the true vertical depth (TVD) of the injection well. When it is not stated

in the literature whether the measured depth (MD) or TVD is given, then the depth is assumed to be TVD. Taking the TVD instead of the MD is important for inclined wells when calculating the pressure and stress at depth. When multiple wells are present at a site and they are not specified in the database, then the reservoir depth is given. An exception is Bouillante, for this site the production depth is given.

• Reservoir depth: The depth of the reservoir, which is the depth at which stimulation/production takes place. This parameter is usually only given when a given site has multiple wells with different depths, because then it is easier to define the depth of the reservoir.

• Open-hole section: The depth range (assumed TVD, unless explicitly mentioned that it is measured depth (MD)) of the uncased part of the well. It is assumed that injection took place at this interval, unless otherwise stated in the literature. For some sites the treatment interval or only the top of the injection interval or casing shoe is given.

• Open-hole height: This is the height of the open-hole section. For some experiments the total perforation interval (sum of all the perforated intervals) is given instead of the open-hole section.

• Borehole diameter: Diameter of the open-hole section. • Temperature: The bottom hole temperature (BHT) or the temperature at a given

depth. • Rock type (reservoir): The type of rocks in the stimulated rock volume/fracture

network. • Young’s modulus (static): An elastic parameter which is a measure of the stiffness of

a rock (derived from static measurements of stress-strain relationship). • Poisson’s ratio (static): An elastic parameter which is a measure of the lateral

expansion relative to longitudinal contraction. An exception is the ratio for The Geysers which is a dynamic value.

• Porosity (reservoir): The percentage of pore volume over the total reservoir volume. It is mentioned whether this applies for virgin reservoir or stimulated reservoir.

• Permeability (reservoir): A measure of the ability of the reservoir rock to transmit fluids.

• Magnitude Sv: Magnitude of the vertical stress (assumed to be the overburden or lithostatic pressure) at the casing shoe (top of the open-hole section). If this depth is unknown, then the stress is given at the bottom hole depth or at another specified depth. This stress value includes the hydrostatic pressure. For some values the error is also given.

• Orientation Sv: Orientation of the vertical stress. Sv is always oriented vertical, because of the assumption that the vertical stress is equal to the overburden.

• Magnitude Sh: Magnitude of the minimum horizontal stress at the casing shoe (top of the open-hole section). If this depth is unknown then the stress is given at the bottom hole depth or at another specified depth. This stress value includes the hydrostatic pressure. For some values the error is also given.

• Orientation Sh: Orientation of the minimum horizontal stress in the reservoir. • Magnitude SH: Magnitude of the maximum horizontal stress at the casing shoe (top

of the open-hole section). If this depth is unknown then the stress is given at the bottom hole depth or at another specified depth. This stress value includes the hydrostatic pressure. For some values the error is also given.

• Orientation SH: Orientation of the maximum horizontal stress in the reservoir or otherwise at a given depth range.

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• SH - Sh: The difference between the maximum horizontal stress and the minimum horizontal stress (differential stress).

• Orientation faults: The predominate orientation (strike and/or dip) of faults in the depth range of the reservoir.

• Length tectonic faults: The total length of all faults mapped on the top structure maps. • Depth tectonic faults: Depth range of the present faults. • Area tectonic faults: The part of the fault which is triggered during stimulation. • Distance natural faults from injection well: The distance between natural faults and

the injection well, which is important in assessing whether faults can be triggered due to stimulation.

• Orientation of fracture network: The strike and/or dip of the seismic cloud formed during stimulation. So the assumption is made that fractures are present at the location of the microseismic events induced during stimulation.

• Length of fracture network: This is assumed to be the length of the seismic cloud. • Width of fracture network: This is assumed to be the width of the seismic cloud. • Area of fracture network: The length of the seismic cloud multiplied by the width of

the seismic cloud. • Volume of fracture network: The area of the fracture network multiplied by the height

of the fracture network. This volume is only a rough estimate of the seismic volume containing the hypocenters of the induced events.

• Orientation of natural fractures: Strike and/or dip of fractures present before stimulation.

• Orientation hydraulically-induced fractures: Strike and/or dip of fractures created due to stimulation.

• Range of magnitudes: Smallest and largest magnitudes of induced events during injection.

• Number of events: The number of induced events during injection. • Magnitude largest event: The maximum local magnitude of all the induced events. • Delay after first injection: The time between beginning of injection and the first

induced event. • Hydrostatic pressure: The pressure exerted by a fluid at equilibrium due to the force

of gravity at the casing shoe/top of injection interval or at a specified depth. • Maximum formation pressure during stimulation: The maximum pressure at the

casing shoe/top of injection interval or at a specified depth. • Start date – end date: The day (and time) at which injection starts and ends.

Sometimes only the start date is given or only the year or month in which the injection/production took place.

• Duration: The duration of injection/production. • Maximum wellhead pressure: The pressure exerted by the fluid at the wellhead. • How much: The amount of volume injected. • Velocity: The maximum or mean flow rate of the fluid (or minimum and maximum

flow rate).

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4.8 Appendix III Non-dimensionless parameters Below, the following parameters are plotted as a function of maximum observed magnitude (Mmax) at a site: Vseism (volume of seismic cloud); Vinj (injected volume); Tinj (duration of injection) and Qmax (maximum injection rate) (Figure III-1). Just as for the dimensionless parameters plotted versus Mmax in Figure 3.1, there is no clear correlation visible for these non-dimensionless parameters. Bayesian least squares fitting to the data points in Figure III-1, resulted in large values for Χ2 and point average (see section 3 for definitions), meaning that these are not good fits to the data. The figures do suggest that injections into granite are more prone to inducing larger events (M > 2.0). Exceptions are the injections at Berlín (into basalt/andesite) and at Paradox Valley (into limestone), which generated large magnitude events (> 2.0). This was also observed in Figure 3.1, since the same parameters were used here.

3 a) b)

c) d)

Figure III-1 Maximum local magnitude plotted as a function of a) Vseism; b) Vinj; c) Tinj and d) Qmax (from database in Appendix I). Sites are separated into those which have a granitic reservoir (blue) and those which have other reservoir rock (green). The dashed line at Mmax of 2.0 represents the assumed division between seismic events which are felt and those which are not felt by humans.

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Merging the ETH and TNO databases TNO and ETH have decided to merge their databases for EGS sites. The ETH database consists of 41 European sites, mostly EGS but also C02 injection sites. The TNO database consists of 19 sites in Europe, US, Australia and South America.

3.1.1 Parameters to include in the merged database After discussion at the General Assembly and the experience obtained in the GEISER program, the parameters that will be included in the database are separated into site factors, operational parameters and seismicity parameters. Site factors are the parameters associated with the site before injection has started, such as the type of the site, depth, temperature, type of ‘reservoir’ rock (rock type, Young’s modulus, Poisson’s ratio), type of bedrock/ overburden, magnitude and orientation of Sv, Sh and SH at a specific depth, stress ratios Sh/Sv, SH/SV, regional tectonic setting, faulting, distance to injection wells of tectonic faults and orientation faults/fractures. Operational parameters are parameters associated with the injection into the subsurface, such as open-hole or perforated sections, well separation, injection duration, injection pressure, injection volume, injection flow rate, formation pressure and bottom hole pressure. Seismicity parameters are parameters associated with the observed (micro)-seismicity at a site, such as background seismicity (PGA, ‘b’ and ‘a’ values from Gutenberg-Richter, tectonic setting from World stress map), network sensitivity (N-felt, N-reported), detection threshold, local or regional network, induced seismicity during injection (‘b’ and ‘a’ values during injection, number of events associated with injection, magnitude of first seismicity, pressure at first seismicity, injected volume at first seismicity), induced seismicity after injection (‘b’ and ‘a’ values after injection, number of events after shut-in), magnitude largest event, type of magnitude and dimensions seismic cloud.

3.1.2 Application to seismic hazard analysis before start of production After merging of the database, statistical evaluation will take place to find key parameters that describe the observed large magnitude events (LME’s) for each site. The predictive properties of these parameters will depend heavily on the number of sites, but also the type of key parameters found. Site factors will be most important for key parameters that describe the seismic hazard before start of production. Operational parameters may play a role in the mitigation of events during production.

3.1.3 Database access via internet A first step in the process of merging the TNO and ETH dbases has been to create a common room in which files and information can be rapidly accessed by both Institutes. In the ETH domain a document management system (DMS) has been dedicated to this dbase. A DMS is a computer system used to track and store electronic documents and/or images of paper documents, which is also capable of keeping track of the different versions created by different users (history tracking). The system has been realized through Agorum core, a free Open-Source Enterprise Content Management system by Agorum Software GmbH from Germany. One of the main features is the Document-Network-Share. With that the documents within the DMS are shown as a normal network share, so it is usable like any other fileserver.

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The Dbase is at the moment a Literature Dbase, organized and based on the geographical distribution of injection sites. It contains about 350 published papers (and a very minor amount of personal comments and unpublished information) related to induced or not-induced microseismicity. It can be considered relatively complete and up to date only for Europe. Whenever available all the documents are related with a URL address, and commented. The Literature Dbase is managed though an EndNote, in which the parameters of interest, as in the previous chapter, are used as key notes, and allow easy search through all the documents.

3.1.4 Future Work In the next year, ETH will, in cooperation with TNO, merge the two databases into one, and make the merged database available for the GEISER partners. TNO will, in cooperation with ETH, perform statistical seismic hazard analyses on the parameters in the merged database to identify key parameters. Also, the existing database will be extended to more non-European geothermal sites.

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