!

!

! DearAuthor:!!

ThankyouforpublishingwithAGU.Yourproofsareattached.Followingtheannotationguidelinesandreturningtheproofsquicklywillallowpromptonlinepostingofthefinalversionofyourpaper.Pleasenotethatrevisedproofsaresentonlyinthecaseof

!WehaveaddedseveralnewfeaturestothejournalWebsitesandadditionalservicesthatwewanttobringtoyourattention.Theonlineversionofyourpaperisenhancedinseveralways,andwillnowinclude altmetric s,” which continua lly trac ks and lin ks to mentio ns innewsoutlets,twitter,blogs,andothersocialmedia.Alinktoscholarlycitationsisalsoprovided.PapersarealsoavailableinaregularPDFandasaReadCubeversion,whichincludesthealtmetricsinformationanddynamiclinkstoreferencesandfigures.

Pleasefollowoursocialmediasites!Severalofourjournalshavesocialmediasites;theseandotherAGUsitesareindexed here.!If!you!have!not!already,!we!encourage!you!to!create!an!ORCID!(Open!Researcher!and!Contributor!ID)!so!that!your!work!can!be!uniquely!tracked.!Your!ORCID!can!be!created!and!linked!through!your!GEMS!account.!!Finally,!your!feedback!is!important!to!us.!!If!you!have!questions!or!comments!regarding!your!AGU!Publications!experience,!including!information!on!production!and!proofs,!please!contact!us!at!publications@agu.org.!!Forquestionsspecifictothispaper,pleasecontactthe!productioneditoratJGRBprod@wiley.com.Wewillalsobecontactingyousoonafteryourarticleispublishedwithanauthorsurvey.Pleasetakeafewminutestorespondtothis

!

onlinesurvey;yourinputisimportantinimprovingtheoveralleditorialandproductionprocess.

Thank!you!again!for!supporting!AGU.!

!

extensivecorrectionsuponauthorrequest.

“

Sincerely,      Brooks  Hanson  Director,  Publications  AGU  bhanson@agu.org    

USING e-ANNOTATION TOOLS FOR ELECTRONIC PROOF CORRECTION

Required software to e-Annotate PDFs: Adobe Acrobat Professional or Adobe Reader (version 8.0

or above). (Note that this document uses screenshots from Adobe Reader DC.)The latest version of Acrobat Reader can be downloaded for free at: http://get.adobe.com/reader/

Once you have Acrobat Reader open on your computer, click on the Comment tab

(right-hand panel or under the Tools menu).

This will open up a ribbon panel at the top of the document. Using a tool will place a comment in the right-hand panel. The tools you will use for annotating your proof are shown below:

1. Replace (Ins) Tool – for replacing text.

Strikes a line through text and opens up a text

box where replacement text can be entered.

How to use it:

Highlight a word or sentence.

Click on .

Type the replacement text into the blue box that

appears.

2. Strikethrough (Del) Tool – for deleting text.

Strikes a red line through text that is to be

deleted.

How to use it:

Highlight a word or sentence.

Click on ..

3. Commenting Tool – for highlighting a section

to be changed to bold or italic or for generalcomments.

How to use it:

Click on .

Type any instructions regarding the text to bealtered into the box that appears.

4. Insert Tool – for inserting missing textat specific points in the text.

Use these 2 tools to highlight the text where a comment is then made.

How to use it:

Click on .

Click at the point in the proof where the comment

should be inserted.

Type the comment into the box that

appears.

Marks an insertion point in the text and

opens up a text box where comments

can be entered.

Click and drag over the text you need to highlight for the comment you will add.

The text will be struck out in red.

Click on .

Click close to the text you just highlighted.

USING e-ANNOTATION TOOLS FOR ELECTRONIC PROOF CORRECTION

For further information on how to annotate proofs, click on the Help menu to reveal a list of further options:

5. Attach File Tool – for inserting large amounts of

text or replacement figures.

Inserts an icon linking to the attached file in the

appropriate place in the text.

How to use it:

Click on .

Click on the proof to where you’d like the attached

file to be linked.

Select the file to be attached from your computer

or network.

Select the colour and type of icon that will appear

in the proof. Click OK.

The attachment appears in the right-hand panel.

6. Add stamp Tool – for approving a proof if no

corrections are required.

Inserts a selected stamp onto an appropriate

place in the proof.

How to use it:

Click on .

Select the stamp you want to use. (The Approved

stamp is usually available directly in the menu that

appears. Others are shown under Dynamic, SignHere, Standard Business).

Fill in any details and then click on the proof

where you’d like the stamp to appear. (Where a

proof is to be approved as it is, this would

normally be on the first page).

7. Drawing Markups Tools – for drawing shapes, lines, and freeform

annotations on proofs and commenting on these marks.

Allows shapes, lines, and freeform annotations to be drawn on proofs and

for comments to be made on these marks.

How to use it:

Click on one of the shapes in the Drawing

Markups section.

Click on the proof at the relevant point and

draw the selected shape with the cursor.

To add a comment to the drawn shape,

right-click on shape and select OpenPop-up Note.

Type any text in the red box that

appears.

Drawing tools available on comment ribbon

THIS IS NOT AN INVOICE JGRB52152

PLEASE COMPLETE THE PUBLICATION FEE CONSENT FORM BELOWAND

RETURN TO THE PRODUCTION EDITOR WITH YOUR PROOF CORRECTIONS

Please return this completed form and direct any questions to the Wiley Journal Production Editor at JGRBprod@wiley.com.

To order OnlineOpen, you must complete the OnlineOpen order form at:https://authorservices.wiley.com/bauthor/onlineopen_order.aspAuthors who select OnlineOpen will be charged the standard OnlineOpen fee for your journal, but excess publication fees will stillapply, if applicable. If your paper has generated excess publication fees, please complete and return the form below inaddition to completing the OnlineOpen order form online (excess fees are billed separately). If you would like to chooseOnlineOpen and you have not already submitted your order online, please do so now.

YOUR ARTICLE DETAILSJournal: Journal of Geophysical Research: Solid EarthArticle: Albano, M., S. Barba, G. Solaro, A. Pepe, C. Bignami, M. Moro, M. Saroli, and S. Stramondo (2017), Aftershocks, groundwaterchanges and postseismic ground displacements related to pore pressure gradients: Insights from the 2012 Emilia-Romagnaearthquake, J. Geophys. Res. Solid Earth, 122, doi:10.1002/2017JB014009.

OnlineOpen: No Words: 7,647 Tables: 1 Figures: 8 Total Publishing Units: 24

Journal Base Fee: $1,000Excess Publishing Units: 0@$125 _______$0Publication Fee Total: USD $1,000

An invoice will be mailed to the address you have provided once your edited article publishes online in its final format. Pleaseinclude on this publication fee form any information that must be included on the invoice.

Publication Fees and Length Guidelines:http://publications.agu.org/author-resource-center/

Frequently Asked Billing Questions:http://onlinelibrary.wiley.com/journal/10.1002/(ISSN)2169-8996/homepage/billing_faqs.pdf

Purchase Order Instructions:Wiley must be listed as the contractor on purchase orders to prevent delay in processing invoices and payments.

Please provide the information requested below.Bill to:

Name: ------------------------------------------------------------------------------------------------------------------------------------------------------------Institution: -------------------------------------------------------------------------------------------------------------------------------------------------------Address:----------------------------------------------------------------------------------------------------------------------------------------------------------Phone: ------------------------------------------------------------- Email: -------------------------------------------------------------------------------------Signature:--------------------------------------------------------- Date: -------------------------------------------------------------------------------------

Aftershocks, groundwater changes and postseismicground displacements related to pore pressuregradients: Insights from the 2012Emilia-Romagna earthquake

Q2Matteo Albano1 , Salvatore Barba1 , Giuseppe Solaro2 , Antonio Pepe2 ,Christian Bignami1 , Marco Moro1, Michele Saroli3,1, and Salvatore Stramondo1

1Istituto Nazionale di Geofisica e Vulcanologia, Rome, Italy, 2Istituto per il Rilevamento Elettromagnetico dell’Ambiente,National Research Council (CNR), Naples, Italy, 3Dipartimento di Ingegneria Civile e Meccanica, Università degli Studi diCassino e del Lazio meridionale, Cassino, Italy

Abstract During the 2012 Emilia-Romagna (Italy) seismic sequence, several time-dependent phenomenaoccurred, such as changes in the groundwater regime and chemistry, liquefaction, and postseismic grounddisplacements. Because time-dependent phenomena require time-dependent physical mechanisms, weinterpreted such events as the result of the poroelastic response of the crust after the main shock. In ourstudy, we performed a two-dimensional poroelastic numerical analysis calibrated with Cosmo-SkyMedinterferometric data and measured piezometric levels in water wells. The simulation results are consistentwith the observed postseismic ground displacement and water level changes. The simulations show thatcrustal volumetric changes induced by poroelastic relaxation and the afterslip along the main shock fault areboth required to reproduce the amplitude (approximately 4 cm) and temporal evolution of the observedpostseismic uplift. Poroelastic relaxation also affects the aftershock distribution. In fact, the aftershocks arecorrelated with the postseismic Coulomb stress evolution. In particular, a considerably higher fraction ofaftershocks occurs when the evolving poroelastic Coulomb stress is positive. These findings highlight theneed to perform calculations that adequately consider the time-dependent poroelastic effect whenmodeling postseismic scenarios, especially for forecasting the temporal and spatial evolution of stresses aftera large earthquake. Failing to do so results in an overestimation of the afterslip and an inaccurate definition ofstress and strain in the postseismic phase.

1. Introduction

On 20 May 2012, anMw 5.9 earthquake struck the Emilia-Romagna region (northern Italy) (Figure F11a), causingsevere damage and loss of life. The main shock and subsequent aftershocks activated two segments of theFerrara-Romagna thrust system, which are the Ferrara thrusts (blue lines in Figure 1a) and the Mirandolathrusts (red lines in Figure 1a). Thousands of aftershocks were registered [Govoni et al., 2014], six of themwithMw > 5, with the largest aftershock (Mw 5.7) occurring on 29 May (Figure 1a). In the days following the mainshock, the rupture propagated eastward and downdip on the Ferrara thrust system [Govoni et al., 2014] andwestward on the adjacent Mirandola thrust, where the Mw 5.7 occurred 9 days later. All aftershocks withMw > 5 occurred within 15 days of the 20 May event, and more than 70% of the earthquakes took placewithin 3 months of the main shock [ISIDe Working Group, 2010]. The moment tensor solutions for strongerevents display a reverse focal mechanism [Scognamiglio et al., 2012].

After the main shock, several postseismic phenomena occurred. The monitored wells close to the epicentralarea registered a sudden increase in water level that was followed by a decay that did not reach the preseis-mic level during the observation period [Marcaccio and Martinelli, 2012]. Some wells ejected a mixture ofwater and sand, and more than 700 liquefaction phenomena were observed [Alessio et al., 2013; Chiniet al., 2015]. Furthermore, a periodic sampling of both liquid and gaseous phases in surficial aquifers estab-lished that deep-seated fluids was mobilized after the main shock and reached surface layers along faultsand fractures, thus altering the geochemical composition of fluids [Italiano et al., 2012; Sciarra et al., 2012].

Postseismic deformation has been partly studied, at least for the Mw 5.7 event that occurred on 29 May.Spanning roughly 1 year after the main shock, differential synthetic aperture radar interferometry (DInSAR)

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 1

PUBLICATIONSJournal of Geophysical Research: Solid Earth

RESEARCH ARTICLE10.1002/2017JB014009

Key Points:• The postseismic phenomena thatoccurred after the 2012 Mw 5.9northern Italy earthquake dependedon the poroelastic response of thecrust

• Poroelastic relaxation is equallyimportant as the afterslip in terms ofthe aftershock distribution

• DInSAR ground deformation data andin-well water level changes were usedto successfully calibrate theporoelastic numerical model

Supporting Information:• Supporting Information S1

Correspondence to:M. Albano,matteo.albano@ingv.it

Citation:Albano, M., S. Barba, G. Solaro, A. Pepe,C. Bignami, M. Moro, M. Saroli, and S.Stramondo (2017), Aftershocks,groundwater changes and postseismicground displacements related to porepressure gradients: Insights from the2012 Emilia-Romagna earthquake,J. Geophys. Res. Solid Earth, 122,doi:10.1002/2017JB014009.

Received 19 JAN 2017Accepted 12 JUN 2017Accepted article online 13 JUN 2017

©2017. American Geophysical Union.All Rights Reserved.

Journal Code Article ID Dispatch: 17.06.17 CE: AMIJ G R B 5 2 1 5 2 No. of Pages: 17 ME:

Revised proofs are sent only in the case of

extensive corrections upon request

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

and GPS time series provided postseismic ground displacement data. Cheloni et al. [2016] highlighted a rapidpostseismic uplift of approximately 2 cm following the Mw 5.7 aftershock on 29 May. These authors assumedthat a transient seismic/aseismic slip located on the Mirandola and Ferrara structures caused the postseismicdisplacement and the aftershock spatial and temporal distributions.

Different time-dependent physical mechanisms may account for one or more postseismic events, e.g., theafterslip [Bürgmann et al., 2002], viscoelastic relaxation [Bürgmann and Dresen, 2008], gravitational rebalan-cing [Albano et al., 2015], and poroelastic rebound [Jónsson et al., 2003]. More than one physical mechanismmay develop after a large earthquake. A poroelastic crustal rebound can explain the observed water levelchanges, the postseismic ground displacements, and the aftershock distribution and decay. Indeed, thecrustal poroelastic rebound has been invoked to explain the location and rate of aftershocks [Bosl and Nur,2002; Shapiro et al., 2003; Antonioli et al., 2005], groundwater modifications [Whitehead et al., 1985; Sneedet al., 2005; Manga and Wang, 2007], and the amplitude and rate of postseismic displacements [Peltzeret al., 1998; Jónsson et al., 2003] after large earthquakes.

Figure 1. (a) Tectonic sketch of the study area and earthquakes ofMw> 2 recorded from May 2012 to May 2014. Circles: earthquakes with magnitudeMw> 2. Stars:earthquakes withMw> 5.0. The red and green stars indicate the events of 20 and 29 May 2012, respectively. The tectonic framework is depicted by the thrusts (solidlines) and back-thrusts (dashed lines) of Ferrara (blue), Mirandola (red), and Pede-Apennines (black) [Boccaletti et al., 2004; Chiarabba et al., 2014]. The purple polygonindicates the footprint of the COSMO-SkyMed (CSK) SAR frame. The dashed black segments indicate the selection limits for earthquakes that are associated with theFerrara thrusts (see Figure 7). The dotted blue and red boxes identify the Ferrara and Mirandola fault planes obtained from the coseismic inversion modeling ofCheloni et al. [2016]. (b) Geological section along the A-A0 profile in Figure 1a [Carminati et al., 2010]. The geologic units are as follows: 1—Quaternary sediments; 2—marine terrigenous deposits (upper to middle Pliocene); 3—evaporitic and terrigenous deposits (upper Miocene to lower Pliocene); 4—marly calcareous and terri-genous levels (Oligocene-Miocene); 5—shallow to deep-water carbonates (Jurassic-Cretaceous) and Triassic evaporites; and 6—crystalline basement.

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 2

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

The crust of the Earth behaves as a biphasic medium comprising of a solid skeleton and voids filled with fluids[Fyfe, 2012]. At seismogenic depths (8–15 km), such fluids are water or water based. Thus, we can assume thatthe crust is a fluid-saturated, biphasic medium.

When an earthquake affects a fluid-saturated medium, the crustal deformation causes a sudden change influid pore pressure with respect to hydrostatic condition. These pore pressures are not in equilibrium, andthe fluid then diffuses from overpressurized regions to under pressurized regions until hydrostatic conditionsare achieved. The amplitude of the pore pressure change and the duration of the diffusion process obey theporoelastic diffusion laws and depend on the elastic and hydraulic properties of the medium and interstitialfluid [Biot, 1941; Rice and Cleary, 1976]. The slow change in pore pressure caused by fluid diffusion modifiesthe effective stresses inside the crust of the Earth, i.e., the stresses that are effective in causing displacements[Terzaghi, 1943], according to equation (1) [Wang, 2000]:

σ0ij ¼ σ ij þ αpδij (1)

where σ0ij and σij are the components of the effective and total stresses (positive when the stresses are

tensile), p is the pore pressure, δij is the Kronecker delta, and α is the Biot -Willis coefficient.

Therefore, further displacements occur at the ground surface and aftershocks occur at depth [Harris, 1998;Bosl and Nur, 2002; Jónsson et al., 2003; LaBonte et al., 2009; Hughes et al., 2010].

Our goal is to investigate the influence of poroelastic rebound on the widespread postseismic phenomenaobserved after the 20 May 2012 Emilia-Romagna earthquake. To achieve this objective, we calculate thestress perturbation and excess pore pressure produced by the fault dislocation associated with the 20 Mayevent using linear poroelasticity. The poroelasticity theory allows us to couple the rock matrix deformationand fluid flow and simultaneously solve for the stress and pore pressure changes. The modeling results areevaluated with regard to the DInSAR ground displacement data obtained for the 20 May event, the measuredcoseismic and postseismic piezometric levels in the wells, and aftershock distribution and decay.

2. DInSAR Analysis and Results

We studied the coseismic and postseismic displacements associated with the 20 May 2012 event using anunpublished data set consisting of 18 synthetic aperture radar (SAR) images acquired on the descendingorbit (side-looking angle of approximately 32°) of the COSMO-SkyMed (CSK) constellation satellites. TheSAR data set, which spans the period from 19 May 2012 to 10 May 2013 (Table S1), allows us to measurethe coseismic and 1 year of the postseismic displacement field. The first postseismic acquisition is dated23 May 2012 (i.e., only 3 days after the main shock), capturing the earliest postseismic displacements. Thefootprint of the CSK data frames is shown in Figure 1a. Although the SAR images partially cover the mainshock area, they represent the best available data set to study the 20May earthquake and related postseismicphenomena. Beginning with the available SAR data, 55 interferograms were generated (Figure S1) andprocessed with a well-established small baseline subset (SBAS) algorithm [Berardino et al., 2002] to retrievea spatially dense map showing the mean displacement velocity of the investigated area and the displace-ment time series for each coherent distributed target (DT) of the scene. The DInSAR interferograms wereflattened using a 90 m Shuttle Radar Topography Mission (SRTM) digital elevation model (DEM) of the area[Rosen et al., 2001]. To mitigate the effects of noise, a complex multilook (ML) operation [Rosen et al., 2000]with 10 looks in both the azimuth and range directions and a noise-filtering operation [Goldstein andWerner, 1998] were independently conducted on each interferogram. The interferograms were unwrappedin space and time by applying the method proposed by Pepe and Lanari [2006] and subsequently invertedusing the SBAS strategy. Residual orbital phase artifacts and topographic mistakes were also corrected[Manzo et al., 2012]. Finally, for each analyzed SAR pixel, the atmospheric phase screen (APS) was estimatedand filtered out from the obtained displacement time series based on the assumption that APSs are possiblycorrelated in space but poorly correlated in time [Ferretti et al., 2001].

The SBAS processing returned 97,902 DT measurement points over an area of approximately 1270 km2

(density of approximately 80 DT/km2) (represented by blue dots in Figure S2). The data from DInSAR areherein presented in terms of mean displacement rate maps and time series. DInSAR can only measure theradar line-of-sight (LOS)-projected component of the surface displacements. We retrieved the ground

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 3

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

displacement pattern by spatially interpolating the displacement values of each DT using the inversedistance to power method.

The coseismic LOS displacement pattern on 23 May 2012 (Figure F22a) shows that the ground moved towardthe satellite (positive values) over an area encompassing the seismic activity in the hanging wall of the mainshock, close to the town of Finale Emilia. A significant component of this motion involves uplift across theInner and Middle Ferrara thrusts (IF and MF in Figure 1a). Although the obtained displacement lobe partiallyimages the extent of the deformed area, it perfectly overlies the coseismic displacement pattern retrieved bythe processing of two Radarsat-2 (RSAT-2) SAR images [Tizzani et al., 2013] (dashed black lines in Figure 2a).However, the CSK coseismic LOS displacement peak at point 1 (Figure 2a) (approximately 16 cm) is lowerthan the value taken from the coseismic displacement field of the Radarsat-2 data at the same point(17.5 cm) (Figure S3). Given that the acquisition geometry for CSK and RSAT-2 is nearly the same, such adifference is explained by considering that the processed Radarsat-2 pairs are dated 30 April 2012 to 17June 2012, and the unwrapped displacement field encompasses approximately 28 days of postseismic uplift(from 20 May to 17 June).

Figure 2. Ground displacement field encompassing the entire seismic sequence from the CSK SAR data along thedescending track (view angle of approximately 32°). (a) Coseismic LOS displacement field associated with the 20 Mayevent (19–23 May). The black dotted lines refer to the coseismic displacement pattern retrieved by the processing of twoRadarsat-2 (RSAT-2) SAR images [Tizzani et al., 2013]. (b) Cumulative postseismic LOS displacement field (23 May 2012 to 10May 2013). The white circles indicate the location of liquefaction phenomena. (c) Mean trend and associated standarddeviation of the displacement time series retrieved from a group of DT close to the red, blue, green, and yellow triangles(indicated with the numbers 1, 2, 3, and 4, respectively) in Figure 2b.

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 4

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

The cumulative postseismic LOS displacements after approximately 1 year following the main shock (10 May2013) (Figure 2b) encompass the entire postseismic phase and show a complex deformation pattern. Thesouthwestern edge of the CSK SAR data frame is affected by a diffused pattern of subsidence that increaseslinearly over time (Figure 2b and point 4 in Figure 2c). Such ground displacements are unrelated to theseismic sequence. Indeed, such an area is affected by anthropogenic subsidence (Figure S2) induced bygroundwater withdrawal [Stramondo et al., 2007; Modoni et al., 2013].

Proceeding to the NE, the area between S. Agostino and Mirabello villages has undergone postseismicground subsidence (negative values in Figure 2b, representing movements away from the sensor). TheLOS displacement time series (point 3 in Figure 2c) show an initial slight postseismic uplift, which is followedby a rapid subsidence that asymptotically decreases over approximately 3 months. Such diffused pattern ofground displacements is ascribable to the widespread liquefaction phenomena observed in the area a fewdays after the main shock [Alessio et al., 2013] (white circles in Figure 2b). Such phenomena (i.e., sand boilsand water leaks from cracks) generally produce ground settlements not only in the affected area but alsowhere the liquefaction phenomena are not recognized [Chini et al., 2015].

Finally, a main lobe of positive LOS displacements (representative of ground movements toward the sensor)with a displacement peak of approximately 3.5 cm is found close to the area that experienced the maximumcoseismic uplift (Figure 2b). In this area, the uplift rate is quite fast given that more than 50% of the recordedpostseismic displacement occurred in the first 2 to 3 months after the earthquake (points 1 and 2 inFigures 2b and 2c). The spatial correspondence between the aftershock epicenters and the coseismic grounddisplacement pattern suggests that the observed postseismic displacements have a tectonic origin [Cheloniet al., 2016].

3. Methods

The observed postseismic uplift has been interpreted as the result of the postseismic poroelastic rebound ofthe crust. Thus, we tested this hypothesis using a fully coupled poroelastic finite element numerical model.Many authors have exploited linear poroelasticity as a useful framework to describe postseismic displace-ments [Peltzer et al., 1998; Jónsson et al., 2003], aftershock evolution over space and time [Bosl and Nur,2002], and the change in well water levels [Nespoli et al., 2016].

3.1. The Poroelastic Model

The three-dimensional coupling between fluid flow and medium deformations belongs to the theory of lin-ear poroelasticity [Biot, 1941; Rice and Cleary, 1976; Wang, 2000]. This theory outlines how the stresses aretransferred from the solid skeleton to the pore fluid, and vice versa.

We exploited the built-in diffusion-structural module of the finite element software package MSCMARC 2015[MSC Software Corporation, 2015], which allowed us to perform fully coupled poroelastic simulations assum-ing realistic fault geometries and spatially variable material properties.

The constitutive laws that relate strains, fluid mass content per unit volume, stresses, and pore pressure are asfollows [Segall, 2005]:

2Gεij ¼ σij � ν1þ ν

σkkδij þ 1� 2νð Þα1þ ν

pδij (2)

Δm ¼ m�m0 ¼ 1� 2νð Þαρ02G 1þ νð Þ ∙ σkk þ 3

Bp

� �(3)

qi ¼ �ρ0κη

∂p∂xj

� ρ0gδij

� �(4)

Equation (2) relates the stresses and strains in a poroelastic medium, where G is the shear modulus; ν is thedrained Poisson ratio; εij and σij are the strain and the stress tensor components, respectively; p is the porepressure; δij is the Kronecker delta and α is the Biot-Willis coefficient.

Equation (3) relates the change in fluid mass per unit volume (Δm) to both the sum of normal stresses (σkk)and pore pressure changes (p). In equation (3), ρ0 is the fluid density, m is the fluid mass content, m0 is the

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 5

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

fluid mass content for a reference state given by the product of the fluid density ρ0 and the volume of thepore fluid per unit volume of solid (i.e., the porosity n), and B is the Skempton coefficient.

Equation (4) is the constitutive law governing the pore fluid diffusion, i.e., the Darcy’s law, where qi is the fluidmass flow rate, κ is the permeability, η is the fluid dynamic viscosity, and g is the gravitational acceleration.Equations (3) and (4) are combined with the mass conservation for the infiltrating pore fluid to obtain thediffusion equation.

The poroelastic formulation implemented into the MARC code assumes that the solid grains constituting themedium are incompressible, and the Skempton (B) and Biot-Willis (α) coefficients can be written as follows[Rice and Cleary, 1976; Wang, 2000]:

B≅Kf

nK þ Kf(5)

α≅1 (6)

where K and Kf are the bulk moduli of the frame and the pore fluid, respectively.

Given the modeling assumptions, a set of eight parameters are required, i.e., the shear modulus G, thedrained frame bulk modulus K (or, equivalently, the drained Young modulus E and Poisson’s ratio ν), the fluidbulk modulus Kf, the porosity n, the permeability k, and the fluid density ρ0 and dynamic viscosity η.

3.2. Finite Element Model Setup

The WNW-ESE alignment of the buried tectonic structures below the Po plain suggests that the fault slipvector is nearly orthogonal to the fault strike line, whereas the along-strike component of displacement isinhibited (i.e., rake ≈ 90°). This hypothesis is confirmed by seismological analyses and both coseismicanalytical and numerical fault solutions (dotted blue and red boxes in Figure 1a) [Bignami et al., 2012;Pezzo et al., 2013; Cheloni et al., 2016]. To take advantage of the symmetry in the displacement, the geologicalsection A-A0 (Figure 1b) is modeled to be parallel to the fault slip vector, i.e., orthogonal to the fault strike(Figure 1a).

Moreover, a 2-D plane strain finite elementmodel (Figure F33) that extends 70 km horizontally and to a depth of20 km is constructed. The mesh is composed of eight-node, isoparametric quadrilateral elements (20,156elements). The elements have a variable size, being coarser with increasing depth; the size ranges fromapproximately 0.35 km at the upper boundary to 0.875 km at the bottom of the model.

The mechanical boundary conditions consist of roller support at the sides and bottom. Hydraulic boundaryconditions consist of atmospheric pressure constraints at the upper edge, and the lower edge is assumedto be impermeable. The model sides are assigned the hydrostatic pore pressure condition to simulate aflowing boundary.

The model geometry is discretized in five strata (Figure 3): a Quaternary layer (layer 1 in Figure 1b), a Pliocenelayer (layers 2 and 3 in Figure 1b), an Oligocene-Miocene layer (layer 4 in Figure 1b), a Jurassic-Cretaceous

Figure 3. Simplified finite element model of the geological cross section A-A0 of Figure 1b. The model grid extends 70 kmhorizontally and to depths of 20 km. The red line represents the 20 May 2012 causative fault according to Cheloni et al.[2016]. The upper (up) and lower (down) red arrows represent the applied edge loads to simulate the coseismic slip. Theblack dashed square indicates the area of contour plots of Figures 4, 6, and 7.

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 6

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

layer (layer 5 in Figure 1b), and a crystalline basement (layer 6 in Figure 1b) [Boccaletti et al., 2004; Carminatiand Vadacca, 2010; Carminati et al., 2010; Tizzani et al., 2013].

The 20 May event causative fault corresponds to the Middle Ferrara thrust (MF in Figures 1a and 1b) [Cheloniet al., 2016]. The fault is modeled using a contact interface [Carminati and Vadacca, 2010; Doglioni et al., 2011]dipping 40° and extending approximately 6 km (red line in Figure 3). Nodes are doubled at the interface; thus,the upper and lower parts of the domain can move relative to each other. The contact interface is assumed tobe impermeable; thus, fluid cannot flow across it. Although fault planes are generally considered permeablesurfaces, observations and experimental data show that as slip increases, faults develop a thick, impermeablegouge layer [Scholz, 1987; Parsons et al., 1999]. Consequently, the modeled fault mostly behaves as a barrierto fluid flow [Piombo et al., 2005].

The slip is simulated by applying edge loads to the top and bottom edges of the fault (up and down arrows inFigure 3). These loads are parallel but have different magnitude and opposite direction.

The selection of appropriate elastic and hydraulic properties of the modeled geomaterials is not straightfor-ward. In fact, local geometric and rheological anisotropies, joint density and orientation, and previous stresspath significantly influences the mechanical and hydraulic properties of rocks at the scale of tens of meters[Carminati et al., 2010].

Model parameters depend on the size of the domain. In our kilometer-scale model, an equivalent continuumapproach is adopted [Sitharam et al., 2001]. It is assumed that each stratum is continuous and isotropic, andwe derive the equivalent parameters from the geomechanical properties of both the intact rock and joints.

The elastic constants were derived from field and literature data. In particular, drained Poisson’s ratios areselected according to literature data [Peltzer et al., 1998; Carminati and Vadacca, 2010; Tizzani et al., 2013].The stiffness of layer 1 (Quaternary) is constrained with experimental data from a sonic log [Maesano andD’Ambrogi, 2017]. Stiffness for layers 2 to 5 are constrained according to an up-to-date seismic velocity modelobtained in the study area after the 2012 seismic sequence [Carannante et al., 2015] (Figure S4). Then,dynamic Youngmoduli (Table T11) are converted to static Youngmoduli according to the empirical relationshipproposed by Brotons et al. [2016] (equation (7)):

Est ¼ 11:531∙ρ�0:457∙E1:251dyn (7)

where Est is the static Young modulus,ρ is the mass density and Edyn is the dynamic Young modulus (FigureS5). The derived static Young moduli (Table 1) are similar to those used in the same area by other numericalmodels [Carminati and Vadacca, 2010; Carafa and Barba, 2011].

Table 1. Elastic, Hydraulic, and State Parameters Adopted in the NumericalQ3 Analysisa

Material property ParameterLayer 1,

QuaternaryLayer 2,Pliocene

Layer 3,Oligocene-Miocene

Layer 4,Jurassic-Cretaceous

Layer 5,Crystalline Basement

Mass densityb ρ(Kg/m3) 2100 2300 2400 2500 2600Drained Poisson’s ratiob ν 0.2 0.28 0.28 0.27 0.26Drained dynamic Young’s modulusc Edyn(Pa) 7.9 × 109 1.3 × 1010 1.8 × 1010 3.7 × 1010 6.7 × 1010

Drained static Young’s modulusd Est(Pa) 4.6 × 109 8.5 × 109 1.2 × 1010 2.9 × 1010 6.1 × 1010

Porositye n 0.2 0.03 0.03 0.03 0.02Permeabilityf k(m2) 5.0 × 10�14 8.0 × 10�14 8.0 × 10�14 8.0 × 10�14 8.0 × 10�15

Fluid bulk modulus Kf (Pa) 2.2 × 109

Fluid dynamic viscosity η(Pa s) 0.001Fluid density ρ0(Kg/m

3) 1000Skempton coefficient B 0.81 0.92 0.89 0.78 0.72Fault friction coefficient μ 0.08

aThese parameters best reproduce the observed pore pressure changes and ground displacements.bDrained Poisson’s ratios obtained from literature data [Peltzer et al., 1998; Carminati and Vadacca, 2010; Tizzani et al., 2013].cDrained dynamic Young’s moduli from Carannante et al. [2015] and Maesano and D’Ambrogi [2017].dStatic Young’s moduli are calculated according to equation (7) [Brotons et al., 2016] as a function of the dynamic Young’s moduli (Edyn) and the rock mass

density (ρ).eMean value from Styles et al. [2014].fValues from Nespoli et al. [2016] (quaternary) and Styles et al. [2014] (other layers).

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 7

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

The hydraulic parameters were selected according to the literature and field data, when available. For theQuaternary layer, an isotropic permeability is assumed because of the large scale of the numerical model.However, according to Lambe andWhitman [1969], the presence of several subhorizontal layers of alternatingaquifers and aquitards considerably reduces the equivalent vertical permeability relative to the equivalenthorizontal permeability. Therefore, assuming that the fluid flow is mainly vertical, we adopted a permeability(Table 1) close to the smallest value from Nespoli et al., [2016].

For the rocky strata, experimental data from diffusivity tests performed at depths between 2.7 and 3.1 kmshow that the permeability ranges between 4 × 10�12 m2 and 3 × 10�15 m2 (Figure S6) and that the porosityranges between 1% and 7% [Laboratorio di monitoraggio Cavone, 2014; Styles et al., 2014]. Such experimentalvalues are assumed to be typical of layers 2, 3, and 4, as these layers have been highly deformed by tectonics.Therefore, the permeability depends more on the degree of fracturing than on the depth. Accordingly, a sin-gle value of permeability for layers 2, 3, and 4 has been selected through a parametric analysis to numericallyreproduce the amplitude and evolution over time of the observed pore pressures and postseismic grounddisplacements, and the other parameters have been held constant. The best value (Table 1) falls inside theexperimental interval (Figure S6) and is common to other numerical models [LaBonte et al., 2009] and litera-ture data for the upper crust [Kuang and Jiao, 2014]. The results for a range of permeability values will bediscussed later.

For layer 5, we assumed a permeability one order of magnitude lower than the above strata to account for thedependency of permeability on depth [Manga et al., 2012; Kuang and Jiao, 2014].

The coupling among the solid and fluid phases is simulated through three steps. In the first step, the initialstress and pore pressure distribution are established by applying a gravity load. The pore pressure distribu-tion at the end of this phase increases linearly with depth according to field measurements [Laboratorio dimonitoraggio Cavone, 2014].

In the second step, the fault responsible for the 20 May event is allowed to slip instantaneously by imposingtwo nonuniform edge loads on the upper and lower faces of the fault (Figure 3). No friction is assumed at thecontact interface at this stage. The magnitudes of the loads (up = 3.25 × 106 N/m2 and down = 2.6 × 106 N/m2) are calibrated following a trial and error procedure to fit the coseismic CSK LOS displacements (Figure 2a)along section A-A0, i.e., the ground displacement data available 3 days after the main shock (23 May 2012).

In the third step, no further loads are applied. The stress changes caused by the coseismic Δp relaxation arecalculated for 2 years after the 2012 main shock. In this phase, a friction law obeying the Coulomb failure cri-terion has been assumed at the contact interface representing the main fault. The friction coefficient (μ) hasbeen parametrically increased, ranging from an unlocked fault (μ = 0) until reaching to a locked fault (μ = ∞),to fit the observed postseismic displacements. During the postseismic phase, the evolution of the fluid porepressures and displacements at the ground has been monitored over time.

3.3. Coulomb Stress Changes

The change in Coulomb stress (ΔCFS) is determined as follows [Cocco and Rice, 2002]:

ΔCFS ¼ Δτ þ μ∙ Δσn þ Δpð Þ ¼ Δτ þ μ∙Δσ0n (8)

where Δτ is the shear stress change (computed in the slip direction), μ is the static friction coefficient [Kinget al., 1994; Harris, 1998], Δσn is the fault-normal total stress change, Δp is the pore pressure change, andΔσ

0n

is the effective fault-normal stress change. The friction coefficient ranges between 0.6 and 0.8 for most rocks[Harris, 1998]. Because crustal rocks are fractured, the lower bound for the friction coefficient is assumed(μ = 0.6) [Piombo et al., 2005], and the Coulomb stresses on preferential fault planes dipping 30° southwestwith respect to the horizontal are calculated according to the mean focal mechanisms of the aftershocksequence [ISIDe Working Group, 2010]. Because the numerical model adopts a positive sign convention fortension and a negative convention for compression, a positive ΔCFS indicates fault weakening.

4. Results

Here we show the results that best reproduce the measured coseismic and postseismic displacements andthe water table changes at wells. The performed parametric analyses will be discussed later. Results arepresented in terms of pore pressures and ground deformations in the coseismic and postseismic phases.

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 8

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

Regarding pore pressures, both positive and negative changes with respect to the hydrostatic values arehereinafter indicated using Δp, and we use “suprahydrostatic” in the case of a pore pressure increase and“subhydrostatic” in the case of a pore pressure decrease.

4.1. Coseismic Δp and Postseismic Evolution

Before the main shock, the pore pressure is hydrostatic. In the coseismic phase, the fault dislocation causesshear and volumetric changes inside the medium. In particular, coseismic volumetric changes develop a Δppattern reaching nearly ±2MPa close to the fault zone (Figure F44a). These Δp values are transient and graduallydissipate in the postseismic phase because of fluid diffusion. Indeed, the Δp values are 4 times lower after10 days (Figure 4b) and 10 times lower after 30 days (Figure 4c). The fluid flow rate (black arrows) is higherin the early days after the earthquake (Figure 4b) and gradually decreases over time (Figure 4c). One year later(Figure 4d), the fluid diffusion ceases and the pore pressures correspond to the hydrostatic values.

The temporal variation in Δp at depth (points 1 and 2 in Figure 4a) shows that subhydrostatic Δp valuesdevelop at point 1 (Figure 4e) after the coseismic slip because of volume dilation. These Δp values

Figure 4. Pore pressure changes and decay with respect to the hydrostatic values. (a) Δp pattern after the simulation of the20 May coseismic slip, (b) 10 days later, (c) 30 days later, and (d) 1 year later. The dashed lines in Figures 4a–4d separateareas with positive and negative Δp. (e) Coseismic Δp peak and postseismic decay at points 1 and 2 (yellow points) inFigures 4a–4d. (f) Coseismic and postseismic Δp values at point 3 compared with the measured water level change at wellFE81 (Figure 1a).

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 9

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

gradually diminish over approximately3–6 months. At point 2, the coseismicvolume compression induces suprahy-drostatic Δp values. These Δp vlauesrapidly decay, reaching subhydrostaticvalues after approximately 70 daysbecause of the arrival of a subhydro-static fluid pulse from the neighboringregions of the domain. SubhydrostaticΔp values gradually return to hydro-static values overapproximately 3 months.

At the surface (point 3 in Figure 4a),coseismic suprahydrostatic Δp valuesdevelop (Figure 4f). However, the pres-sure peak is lower than that at depth(e.g., at point 2). Hydrostatic pore

pressures are rapidly restored because the selected point is very close to the ground surface; thus, it isconstrained by atmospheric pressure. Moreover, the drainage path is very short.

The modeled coseismic and postseismic Δp values are compared with the measured water level change atseveral wells close to the earthquake area. Indeed, the Quaternary aquifer response to the Emilia seismicevent was captured by the monitoring network of the regional agency for environmental protection(Advanced Research Projects Agency) [Marcaccio and Martinelli, 2012]. In particular, three wells (FE80, FE81,and M080 in Figure 1a) recorded a significant water level change after the 20 May event. The water levelexperienced a sudden increase between 0.2 and 1.6 m, which was followed by a gradual decay that didnot recover to the pre-earthquake water level in the observed time interval. The highest peak was observedfor the deepest well (M080, depth = 300 m), which was followed by a slower decay with respect to theshallower wells (FE80, depth = 40 m; FE81, depth = 40 m). A smaller peak was also observed after the May29 event [Marcaccio and Martinelli, 2012].

The comparison has been performed at the FE81 well, as this well is close to the modeled cross section. WellsM080 and FE80 were not considered because earthquake-induced Δp values have a great horizontal variation[Nespoli et al., 2016], and these wells are too far from the modeled 2-D section. The modeled Δp values atpoint 3 were converted into the equivalent water level change (1 MPa corresponds to a water column heightof 101.97 m at 4°C) and compared with the measured water level change (Figure 4f). The observed coseismicpeak and postseismic decay are well reproduced by the model, thus validating the Δp pattern shown inFigure 4a.

4.2. Coseismic and Postseismic Deformations

In the coseismic phase, the modeled coseismic slip (Figure F55) produces a deformation pattern (Figure F66a) thatmainly involves a localized area of the medium above the dislocated fault. The displacement pattern andvectors, with a maximum amplitude of approximately 30 cm, highlight the slip along the causative faultand agree with the solution proposed by Tizzani et al. [2013] that was obtained from a similar 2-D finiteelement model, which assumed complete dryness. Our model is completely saturated; thus, the imposedcoseismic slip results in the Δp pattern shown in Figure 4a.

In the postseismic phase, the gradual dissipation of the coseismic Δp alters the stresses inside the medium,developing further deformations.

Indeed, the modeled postseismic deformations 2 years after the 20 May event (Figure 6b) show a furtheruplift close to the area already affected by the coseismic inflation, which is consistent with the observedpostseismic uplift from DInSAR (Figure 2b) and analogous postseismic scenarios [Fattahi et al., 2015;Cheloni et al., 2016].

The modeled coseismic and postseismic deformations are compared with the DInSAR-derived LOS displa-cements along section A-A0. In particular, the first (23 May 2012) and last (10 May 2013) available

Figure 5. Modeled coseismic (red line) and postseismic (blue lines) slip.The modeled fault is unlocked in the coseismic phase (i.e., μ = 0), whilethe friction coefficient is parametrically increased in the postseismicphase to fit the observed ground displacements from DInSAR. The plottedpostseismic slip is obtained for μ = 0.08 (Table 1).

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 10

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

displacement profiles are compared with the modeled LOS displacement profiles taken 3 and 355 daysafter the main shock, respectively (Figure 6c). The agreement between the modeled (blue line) andmeasured (blue triangles) LOS displacements 3 days after the main shock (Figure 5c) is very good; theroot-mean-square error (RMSE) is less than 1 cm (RMSE = 0.64 cm) and comparable with the modelingresults of Tizzani et al. [2013]. The best fit modeled LOS displacements 355 days after the main shock(red line) (with μ = 0.02) show a satisfactory agreement with the measured DInSAR movements (redcircles) (RMSE = 1.01 cm), confirming the effectiveness of the model geometry, parameters, andboundary conditions.

Concerning the temporal evolution of the LOS displacements, the modeled time series (red line in Figure 6d)at point 1 (Figure 6a) show that more than 50% of the observed postseismic uplift occurred in the first2–3 months after the event. The modeled LOS displacement trend follows the mean trend of the DInSARdisplacements (red circles in Figure 6d) measured at some DT close to the point 1 in Figure 2 both in termsof shape and amplitude (RMSE = 0.36 cm).

5. Discussion

The performed simulations highlight that the postseismic observations following the 20 May event, i.e., theobserved rates and amplitudes of ground displacements and water table fluctuations, are compatible withporoelastic rebound at seismogenic depths.

In our modeling, the postseismic rebalancing of pore water pressures is the only important driving mechan-ism for the postseismic ground displacement (Figure 6b). In fact, after the coseismic slip, the pore pressureincreases in compressed areas and decreases in dilated areas (Figure 4a). The coseismic pore pressure gradi-ent triggers groundwater flow from regions of higher pore pressures, which further compress in the postseis-mic phase, to regions of lower pore pressures, which further dilate (Figures 4b and S7). Hence, coseismic Δpvalues gradually dissipate.

Figure 6. Coseismic and postseismic deformation patterns and displacement profiles at the surface. (a) Coseismic deforma-tion pattern after the simulated 20 May 2012 event. (b) Cumulative postseismic deformation pattern two years after theevent resulting from the rebalancing of fluid overpressures. (c) Comparison between the measured and modeled LOSdisplacement profiles along the A-A0 section at 3 days (blue line and symbols) and 355 days (red line and symbols) after themain shock. (d) Comparison between themodeled LOS displacement time series at point 1 in Figure 6a and themean trendof DInSAR LOS displacements for a group of DT close to the red triangle in Figures 2a and 2b.

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 11

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

The time required for Δp to dissipate and the magnitude and temporal evolution of the modeled grounddisplacements strongly depend on the hydraulic properties of the medium. Indeed, a reduction in perme-ability of 1 order of magnitude in Table 1 (keeping the other parameters fixed) is associated with a slowdissipation of Δp (dash-dotted curves in Figure S8a). Consequently, postseismic ground displacementsincrease slowly and the observed LOS displacement trend tends to be underestimated (green curve inFigure S8b).

In contrast, increase in permeability of 1 order of magnitude in Table 1 is associated with a fast dissipation ofΔp (dashed lines in Figure S8a) resulting in overestimation of the observed LOS displacement trend in the firstmonths after the main shock (blue curve in Figure S8b).

Most of the diffusion process occurs in layers 2, 3, and 4. Indeed, further parametric analyses (not presented)show that the permeability of layer 5 has a negligible influence on the amplitude and temporal evolution ofΔp and postseismic displacement, while the permeability of layer 1 affects the modeled pore pressure decayonly at the FE81 well (Figure 4f).

The poroelastic compression/dilation of the medium causes the main shock causative fault to further slip inthe postseismic phase (Figure 5). In particular, the amount of afterslip (hence, the amount of postseismicdisplacement) depends on the adopted residual friction at the contact interface. Parametric analyses withdifferent friction coefficients show that the afterslip decreases by increasing the friction coefficient and thatthe best fit solution (μ = 0.08, Figures 6c 6d and S9) corresponds to a mean postseismic slip of approximately3.5 cm. Moreover, the parametric analysis shows that assuming an unlocked fault overestimates the postseis-mic displacements (Figure S9). Conversely, assuming a locked fault underestimates the postseismic displace-ments. Thus, the volumetric strain of the medium is not sufficient to reproduce the observed postseismicground displacements, although the coseismic pore pressures and effective stresses are overestimateddue to the assumption of incompressible grains.

The modeling assumptions and the plane-strain geometry of our modeling do not allow for a proper evalua-tion of the fault afterslip; thus, the contribution of afterslip and volumetric strain to the postseismic grounddisplacements could be different. A 3-D model would certainly produce more accurate results, althoughthe relative importance of afterslip and volumetric strain would be similar. However, this result indicates thatneglecting the poroelastic rebound mechanism, which sometimes occurs in inversion procedures, results inan overestimate of the afterslip. In general, the poroelastic contribution to ground displacement must beproperly quantified in postseismic and interseismic studies.

Finally, the role of pore fluid flow in triggering aftershocks is investigated by calculating the ΔCFS in thecoseismic and postseismic phases and comparing these with the aftershock distribution and rate. In particu-lar, we sought to ascertain whether there are regions where the ΔCFS increases in the postseismic periodbecause of fluid flow in the crust [Piombo et al., 2005].

The coseismic ΔCFS (Figure F77a) highlights the areas that are strengthened (negative ΔCFS) and the areas thatare weakened (positive ΔCFS) after the main shock. The computed ΔCFS pattern is similar to other solutionsproposed by different authors in the same area [Ganas et al., 2012; Pezzo et al., 2013; Bonini et al., 2014].

In the postseismic phase, coseismic Δp values decrease, leading to a ΔCFS change (equation (8)). In particular,at 7 days after the main shock (Figure 7b), the positive lobes to the left of the fault expand with respect to thecoseismic at the expense of the top left negative lobe. At approximately 70 days after the main shock(Figure 7c), the positive lobes extend farther. Some areas where the ΔCFS was initially negative (i.e., earth-quakes inhibited), now show positive values (i.e., earthquakes facilitated). Two years later (Figure 7d), ΔCFSundergoes minor changes with respect to the pattern shown in Figure 7c because the diffusion processterminates after approximately 3–6 months following the main shock (Figure 4c).

The ΔCFS evolution is shown for several points located near the fault trace (Figure 7a). At point 1 (Figure 7e), the coseismic ΔCFS is positive because the increase in coseismic shear stresses (Δτ) prevails with respectto the resistance increase due to the negative (i.e., compressive) effective normal stresses (Δσ

0). In the post-

seismic phase, ΔCFS further increases as the compressive Δσ0gradually decreases because the negative Δp

values increase from subhydrostatic to hydrostatic values. At point 2 (Figure 7f), the coseismic ΔCFS isnegative because the Δτ is negative, meaning that the area is initially strengthened. However, in the post-seismic, the area is gradually weakened because of fluid diffusion. Indeed, the dissipation of Δp gradually

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 12

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

increases the positive (i.e., tensile) Δσ0; thus, ΔCFS increases until reaching positive values. At point 3

(Figure 7g), the coseismic ΔCFS is negative because both the Δτ and Δσ0values are negative. In the

postseismic phase, ΔCFS further decreases because the Δp values decrease from suprahydrostatic valuesto the hydrostatic value.

The modeling results show that only the ΔCFS changes due to poroelastic relaxation are statistically signifi-cant, reaching approximately ±1 MPa close to the fault trace (Figure S10). In particular, the areas with coseis-mic subhydrostatic Δp experience an increase in the postseismic ΔCFS, while the areas with coseismicsuprahydrostatic Δp experience a postseismic ΔCFS reduction.

It is expected that aftershocks occur preferentially in regions characterized by a postseismic ΔCFS increase.Indeed, a spatial correlation is found between ΔCFS and aftershocks belonging to the Ferrara thrusts (i.e.,the aftershocks bordered by the dashed black lines in Figure 1a) [ISIDe Working Group, 2010; Govoni et al.,2014]. A higher percentage of aftershocks occurred in positive ΔCFS areas 2 years after the earthquake(approximately 80%) (Figure 7d) relative to the coseismic ΔCFS (approximately 70%) (Figure 7a).

Figure 7. Model of stress changes (ΔCFS) and earthquakes at different times (see Figure 3 for the location of the model). (a) Coseismic ΔCFS pattern and aftershockson 20 May 2012. Coseismic plus postseismic ΔCFS model and aftershocks at (b) 7 days, (c) 3 months, and (d) 2 years from the 20 May event. The white dots representthe aftershock locations as selected in Figure 1. The dashed lines in Figures 7a–7d separate areas with positive and negative ΔCFS. Evolution over time of ΔCFS, Δp,Δτ, and Δσ0 for points (e) 1, (f) 2, and (g) 3 located in Figure 7a.

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 13

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

Moreover, we calculated the percentage of aftershocks that originated in positive and negative ΔCFS areasat 5, 10, 20, 40, 80 160, 320, and 730 days after the main shock. The results (Figure F88) reveal that thepercentage of aftershocks in positive ΔCFS areas gradually increases at the expense of aftershocks innegative ΔCFS areas. The number of aftershocks rapidly increases soon after the earthquake, accordingto the rapid increase in ΔCFS in the postseismic phase, and attains a constant percentage afterapproximately 3–6 months.

Aftershocks occurring in negative ΔCFS areas rapidly decrease in the first 3–6 months, reaching a percentageof approximately 20%, then slightly increase over the subsequent 570 days. Poroelastic effects do not causethis increase because the diffusion process terminates after approximately 3–6 months (Figure 4c).

The presence of aftershocks in negative ΔCFS areas is partly due to the model approximations or unmodeledphenomena. In fact, the Δp pattern is calculated based on only the 20 May event because the largest after-shocks are not included in our calculations of the pore pressures and stress field. Moreover, a thrust dippingof 30° is assumed. Nonetheless, the anomalous aftershocks can be explained with a different rupturemechanism because small normal faulting aftershocks can occur in a thrust environment [Kato et al., 2011;Bonini et al., 2014]. Finally, the heterogeneities of the material properties significantly affect the ΔCFS distri-bution because the fluid flow is controlled by intrinsic permeability that can be anisotropic and heteroge-neous in the crust.

A more accurate simulation accounting for material heterogeneity, different rupture mechanisms, andaftershock dependency of pore pressure will lead to more accurate results. However, we believe that thefirst-order findings of this work are generally valid.

Our findings agree with similar studies performed in different areas of the world [e.g., Steacy, 2005, andreferences therein] and using different approaches [e.g., Cocco and Rice, 2002; Piombo et al. 2005]. Indeed,the observed water level changes are related to pore fluid diffusion, as observed for the South Iceland seismicevents in 2000 [Jónsson et al., 2003]. Most of aftershocks occur where the postseismic Coulomb stressincreases because of pore fluid diffusion, as in the case of the earthquake sequences of Dobi [Noir et al.,1997] and Landers [Bosl and Nur, 2002]. Moreover, poroelastic rebound has been proposed to explainpostseismic ground deformations [Peltzer et al., 1998; Jónsson et al., 2003].

Within the same area and same earthquake sequence, our interpretation agrees with published studiesdemonstrating that changes in the water chemistry in shallow and deep wells [Marcaccio and Martinelli,2012] are connected to the poroelastic response of the crust to coseismic deformation (Figure 4f).However, our interpretation disagrees with some hypotheses that the postseismic ground displacement isonly ascribed to the afterslip [e.g., Cheloni et al., 2016] because pore pressure is equally important.

The connection between poroelastic stress changes and postseismic phenomena indicates that achievingbetter knowledge of the dynamics of the Earth’s crust is possible. For example, adequately detailed poroelas-tic models can be implemented and studied to forecast the temporal and spatial evolution of stresses after alarge earthquake. Such knowledge can help manage emergencies during seismic sequences of both naturaland anthropogenic earthquakes.

Figure 8. Percentage of aftershocks occurring in areas with positive and negative ΔCFS, calculated at 5, 10, 20, 40, 80 160,320, and 730 days after the main shock. The blue dashed line shows the cumulated number of Q4aftershocks.

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 14

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

6. Conclusions

The postseismic ground displacement after the 20 May 2012Mw 5.9 earthquake in Emilia-Romagna (Italy) andthe possible causes of the postseismic phenomena are investigated. For this purpose, a poroelastic finiteelement model calibrated using Cosmo-SkyMed SAR data and measured piezometric levels in water wellsis used [Marcaccio and Martinelli, 2012].

We found that the poroelastic rebound of the deeper crust drives the postseismic displacement following themain shock. Volumetric changes in the crust and the fault afterslip are equally important for determining theamplitude and temporal evolution of the observed postseismic displacement. Such a poroelastic model alsopredicts the coseismic and postseismic water level changes satisfactorily.

The aftershocks recorded during the Emilia-Romagna sequence correlate spatially and temporally with theΔCFS in the postseismic phase, according to the poroelastic diffusion process. Indeed, considerably moreaftershocks occur in areas where the evolving poroelastic ΔCFS (rather than the elastic component) is positiveat the time of the aftershock.

These findings highlight the need to perform calculations that adequately account for the time-dependentporoelastic effect when modeling postseismic and interseismic scenarios. Failing to do so leads to anoverestimation of the afterslip during the postseismic phase and an inaccurate definition of the stress andstrain fields during the interseismic phase.

Because permeability controls the direction and rate of fluid diffusion in porous media, more accurateknowledge of the permeability structure in and around fault zones allows for better prediction of the rateand amplitude of poroelastic stress change due to pore fluid flow.

ReferencesAlbano, M., S. Barba, M. Saroli, M. Moro, F. Malvarosa, M. Costantini, C. Bignami, and S. Stramondo (2015), Gravity-driven postseismic

deformation following the Mw 6.3 2009 L’Aquila (Italy) earthquake, Sci. Rep., 5, 16558, doi:10.1038/srep16558.Alessio, G., et al. (2013), Liquefaction phenomena associated with the Emilia earthquake sequence of May–June 2012 (Northern Italy), Nat.

Hazards Earth Syst. Sci., 13(4), 935–947, doi:10.5194/nhess-13-935-2013.Antonioli, A., D. Piccinini, L. Chiaraluce, and M. Cocco (2005), Fluid flow and seismicity pattern: Evidence from the 1997 Umbria-Marche

(central Italy) seismic sequence, Geophys. Res. Lett., 32, L10311, doi:10.1029/2004GL022256.Berardino, P., G. Fornaro, R. Lanari, and E. Sansosti (2002), A new algorithm for surface deformation monitoring based on small baseline

differential SAR interferograms, IEEE Trans. Geosci. Remote Sens., 40(11), 2375–2383, doi:10.1109/TGRS.2002.803792.Bignami, C., et al. (2012), Coseismic deformation pattern of the Emilia 2012 seismic sequence imaged by Radarsat-1 interferometry, Ann.

Geophys., 55, 789–795, doi:10.4401/ag-6157.Biot, M. A. (1941), General theory of three-dimensional consolidation, J. Appl. Phys., 12(2), 155, doi:10.1063/1.1712886.Boccaletti, M., M. Bonini, G. Corti, P. Gasperini, L. Martelli, L. Piccardi, P. Severi, and G. Vannucci (2004), Carta sismo-tettonica della Regione

Emilia-Romagna, scala 1: 250.000 e note illustrative.Bonini, L., G. Toscani, and S. Seno (2014), Three-dimensional segmentation and different rupture behavior during the 2012 Emilia seismic

sequence (Northern Italy), Tectonophysics, 630(C), 33–42, doi:10.1016/j.tecto.2014.05.006.Bosl, W. J., and A. Nur (2002), Aftershocks and pore fluid diffusion following the 1992 Landers earthquake, J. Geophys. Res., 107(B12), 2366,

doi:10.1029/2001JB000155.Brotons, V., R. Tomás, S. Ivorra, A. Grediaga, J. Martínez-Martínez, D. Benavente, and M. Gómez-Heras (2016), Improved correlation

between the static and dynamic elastic modulus of different types of rocks, Mater. Struct., 49(8), 3021–3037, doi:10.1617/s11527-015-0702-7.

Bürgmann, R., and G. Dresen (2008), Rheology of the lower crust and upper mantle: Evidence from rock mechanics, geodesy, and fieldobservations, Annu. Rev. Earth Planet. Sci., 36(1), 531–567, doi:10.1146/annurev.earth.36.031207.124326.

Bürgmann, R., S. Ergintav, P. Segall, E. H. Hearn, S. McClusky, R. E. Reilinger, H. Woith, and J. Zschau (2002), Time-dependent distributedafterslip on and deep below the Izmit earthquake rupture, Bull. Seismol. Soc. Am., 92(1), 126–137, doi:10.1785/0120000833.

Carafa, M. M. C., and S. Barba (2011), Determining rheology from deformation data: The case of central Italy, Tectonics, 30, TC2003,doi:10.1029/2010TC002680.

Carannante, S., A. Argnani, M. Massa, E. D’Alema, S. Lovati, M. Moretti, M. Cattaneo, and P. Augliera (2015), The May 20 (MW 6.1) and 29 (MW

6.0), 2012, Emilia (Po Plain, northern Italy) earthquakes: New seismotectonic implications from subsurface geology and high-qualityhypocenter location, Tectonophysics, 1234, 1–17, doi:10.1016/j.tecto.2015.05.015.

Carminati, E., and L. Vadacca (2010), Two- and three-dimensional numerical simulations of the stress field at the thrust front of the NorthernApennines, Italy, J. Geophys. Res., 115, B12425, doi:10.1029/2010JB007870.

Carminati, E., D. Scrocca, and C. Doglioni (2010), Compaction-induced stress variations with depth in an active anticline: Northern Apennines,Italy, J. Geophys. Res., 115, B02401, doi:10.1029/2009JB006395.

Cheloni, D., R. Giuliani, N. D’Agostino, M. Mattone, M. Bonano, G. Fornaro, R. Lanari, D. Reale, and S. Atzori (2016), New insights into faultactivation and stress transfer between en echelon thrusts: The 2012 Emilia, Northern Italy, earthquake sequence, J. Geophys. Res. SolidEarth, 121, 4742–4766, doi:10.1002/2016JB012823.

Chiarabba, C., P. De Gori, L. Improta, F. P. Lucente, M. Moretti, A. Govoni, M. Di Bona, L. Margheriti, A. Marchetti, and A. Nardi (2014), Frontalcompression along the Apennines thrust system: The Emilia 2012 example from seismicity to crustal structure, J. Geodyn., 82, 98–109,doi:10.1016/j.jog.2014.09.003.

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 15

AcknowledgmentsWe thank the Associate Editor and thetwo anonymous reviewers for theircomments and suggestions that helpedto improve this article. We thank theItalian Space Agency for providing theCSK data. Both the data and results ofthe DInSAR processing and of the finiteelement modeling are available fromthe corresponding author upon request(matteo.albano@ingv.it). This researchwas supported by the FP7 projectAPhoRISM (code: D88C13000990006).

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

Chini, M., et al. (2015), Coseismic liquefaction phenomenon analysis by COSMO-SkyMed: 2012 Emilia (Italy) earthquake, Int. J. Appl. Earth Obs.Geoinf., 39, 65–78, doi:10.1016/j.jag.2015.02.008.

Cocco, M., and J. R. Rice (2002), Pore pressure and poroelasticity effects in Coulomb stress analysis of earthquake interactions, J. Geophys.Res., 107(B2), 2030, doi:10.1029/2000JB000138.

Doglioni, C., S. Barba, E. Carminati, and F. Riguzzi (2011), Role of the brittle-ductile transition on fault activation, Phys. Earth Planet. Inter., 184,160–171, doi:10.1016/j.pepi.2010.11.005.

Fattahi, H., F. Amelung, E. Chaussard, and S. Wdowinski (2015), Coseismic and postseismic deformation due to the 2007 M5.5 Ghazabandfault earthquake, Balochistan, Pakistan, Geophys. Res. Lett., 42, 3305–3312, doi:10.1002/2015GL063686.

Ferretti, A., C. Prati, and F. Rocca (2001), Permanent scatterers in SAR interferometry, IEEE Trans. Geosci. Remote Sens., 39(1), 8–20, doi:10.1109/36.898661.

Q5Fyfe, W. S. (2012) Fluids in the Earth’s Crust: Their Significance in Metamorphic, Tectonic and Chemical Transport Process, Dev. Geochem., ElsevierScience. Q6

Ganas, A., Z. Roumelioti, and K. Chousianitis (2012), Static stress transfer from the may 20, 2012, M 6.1 Emilia-Romagna (northern Italy)earthquake using a co-seismic slip distribution model, Ann. Geophys., 55(4), 655–662, doi:10.4401/ag-6176.

Goldstein, R. M., and C. L. Werner (1998), Radar interferogram filtering for geophysical applications, Geophys. Res. Lett., 25(21), 4035–4038,doi:10.1029/1998GL900033.

Govoni, A., et al. (2014), The 2012 Emilia seismic sequence (Northern Italy): Imaging the thrust fault system by accurate aftershock location,Tectonophysics, 622, 44–55, doi:10.1016/j.tecto.2014.02.013.

Harris, R. A. (1998), Introduction to special section: Stress triggers, stress shadows, and implications for seismic hazard, J. Geophys. Res.,103(B10), 24,347–24,358, doi:10.1029/98JB01576.

Hughes, K. L. H., T. Masterlark, and W. D. Mooney (2010), Poroelastic stress-triggering of the 2005 M8.7 Nias earthquake by the 2004 M9.2Sumatra-Andaman earthquake, Earth Planet. Sci. Lett., 293(3–4), 289–299, doi:10.1016/j.epsl.2010.02.043.

ISIDe Working Group (2010), Italian seismological instrumental and parametric database, INGV. [Available at http://iside.rm.ingv.it/iside/standard/index.jsp (Accessed 16 March 2015).]

Italiano, F., M. Liotta, M. Martelli, G. Martinelli, R. Petrini, A. Riggio, A. L. Rizzo, F. Slejko, and B. Stenni (2012), Geochemical features and effectson deep-seated fluids during the May–June 2012 southern Po Valley seismic sequence, Ann. Geophys., 55(4), 815–821, doi:10.4401/ag-6151.

Jónsson, S., P. Segall, R. Pedersen, G. Björnsson, and G. Bjornsson (2003), Post-earthquake ground movements correlated to pore-pressuretransients, Nature, 424(6945), 179–183, doi:10.1038/nature01776.

Kato, A., S. Sakai, and K. Obara (2011), A normal-faulting seismic sequence triggered by the 2011 off the Pacific coast of Tohoku Earthquake:Wholesale stress regime changes in the upper plate, Earth, Planets Sp., 63(7), 745–748, doi:10.5047/eps.2011.06.014.

King, G., C. P. King, R. S. Stein, and J. Lin (1994), Static stress changes and the triggering of earthquakes, Bull. Seismol. Soc. Am., 84(3), 935–953,doi:10.1016/0148-9062(95)94484-2.

Kuang, X., and J. J. Jiao (2014), An integrated permeability-depth model for Earth’s crust, Geophys. Res. Lett., 41, 7539–7545, doi:10.1002/2014GL061999.

LaBonte, A. L., K. M. Brown, and Y. Fialko (2009), Hydrologic detection and finite element modeling of a slow slip event in the Costa Rica prismtoe, J. Geophys. Res., 114, B00A02, doi:10.1029/2008JB005806.

Laboratorio di monitoraggio Cavone (2014), Risultati del Programma di prova di interferenza/iniettività.Q7Lambe, T. W., and R. V. Whitman (1969) Soil Mechanics, John Wiley, New York.

Maesano, F. E., and C. D’Ambrogi (2017), Vel-IO 3D: A tool for 3D velocity model construction, optimization and time-depth conversion in 3Dgeological modeling workflow, Comput. Geosci., 99, 171–182, doi:10.1016/j.cageo.2016.11.013.

Manga, M., and C. Y. Wang (2007), Earthquake hydrology, Treatise Geophys., 4, 293–320, doi:10.1016/B978-044452748-6.00074-2.Manga, M., I. Beresnev, E. E. Brodsky, J. E. Elkhoury, D. Elsworth, S. E. Ingebritsen, D. C. Mays, and C.-Y. Wang (2012), Changes in

permeability caused by transient stresses: Field observations, experiments, and mechanisms, Rev. Geophys., 50, RG2004, doi:10.1029/2011RG000382.

Manzo, M., Y. Fialko, F. Casu, A. Pepe, and R. Lanari (2012), A quantitative assessment of DInSAR measurements of Interseismic deformation:The Southern San Andreas Fault case study, Pure Appl. Geophys., 169(8), 1463–1482, doi:10.1007/s00024-011-0403-2.

Marcaccio, M., and G. Martinelli (2012), Effects on the groundwater levels of the May–June 2012 Emilia seismic sequence, Ann. Geophys.,55(4), 811–814, doi:10.4401/ag-6139.

Modoni, G., G. Darini, R. L. Spacagna, M. Saroli, G. Russo, and P. Croce (2013), Spatial analysis of land subsidence induced by groundwaterwithdrawal, Eng. Geol., 167, 59–71, doi:10.1016/j.enggeo.2013.10.014.

MSC Software Corporation (2015), Marc 2015 volume A:: Theory and user information. [Available at http://www.mscsoftware.com/it/product/marc (Accessed 16 March 2015).]

Nespoli, M., M. Todesco, E. Serpelloni, M. E. Belardinelli, M. Bonafede, M. Marcaccio, A. P. Rinaldi, L. Anderlini, and A. Gualandi (2016),Modeling earthquake effects on groundwater levels: Evidences from the 2012 Emilia earthquake (Italy), Geofluids, 16(3), 452–463,doi:10.1111/gfl.12165.

Noir, J., E. Jacques, S. Békri, P. M. Adler, P. Tapponnier, and G. C. P. King (1997), Fluid flow triggered migration of events in the 1989 Dobiearthquake sequence of Central Afar, Geophys. Res. Lett., 24(18), 2335–2338, doi:10.1029/97GL02182.

Parsons, T., R. S. Stein, R. W. Simpson, and P. A. Reasenberg (1999), Stress sensitivity of fault seismicity: A comparison between limited-offsetoblique and major strike-slip faults, J. Geophys. Res., 104(B9), 20,183–20,202, doi:10.1029/1999JB900056.

Peltzer, G., P. Rosen, F. Rogez, and K. Hudnut (1998), Poroelastic rebound along the Landers 1992 earthquake surface rupture, J. Geophys. Res.,103(B12), 30,131–30,145, doi:10.1029/98JB02302.

Pepe, A., and R. Lanari (2006), On the extension of the minimum Cost flow algorithm for phase unwrapping of multitemporal differential SARinterferograms, IEEE Trans. Geosci. Remote Sens., 44(9), 2374–2383, doi:10.1109/TGRS.2006.873207.

Pezzo, G., et al. (2013), Coseismic deformation and source modeling of the May 2012 Emilia (Northern Italy) earthquakes, Seismol. Res. Lett.,84(4), 645–655, doi:10.1785/0220120171.

Piombo, A., G. Martinelli, and M. Dragoni (2005), Post-seismic fluid flow and Coulomb stress changes in a poroelastic medium, Geophys. J. Int.,162(2), 507–515, doi:10.1111/j.1365-246X.2005.02673.x.

Rice, J. R., and M. P. Cleary (1976), Some basic stress-difusion solutions for fluid saturated elastic porous media with compressibleconstituents, Rev. Geophys., 14(2), 227–241, doi:10.1029/RG014i002p00227.

Rosen, P. A., S. Hensley, I. R. Joughin, F. K. Li, S. N. Madsen, E. Rodriguez, and R. M. Goldstein (2000), Synthetic aperture radar interferometry,Proc. IEEE, 88(3), 333–382, doi:10.1109/5.838084.

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 16

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

Rosen, P. A., S. Hensley, E. Gurrola, F. Rogez, S. Chan, J. Martin, and E. Rodriguez (2001), SRTM C-band topographic data: Quality assessmentsand calibration activities, in IGARSS 2001. Scanning the Present and Resolving the Future. Proceedings. IEEE 2001 International Geoscience andRemote Sensing Symposium (Cat. No.01CH37217), vol. 2, pp. 739–741, IEEE, Piscataway, NJ.

Scholz, C. H. (1987), Wear and gouge formation in brittle faulting, Geology, 15(6), 493, doi:10.1130/0091-7613(1987)15<493:WAGFIB>2.0.CO;2.

Sciarra, A., B. Cantucci, M. Buttinelli, G. Galli, M. Nazzari, L. Pizzino, and F. Quattrocchi (2012), Soil-gas survey of liquefaction and collapsedcaves during the Emilia seismic sequence, Ann. Geophys., 55(4), 803–809, doi:10.4401/ag-6122.

Scognamiglio, L., et al. (2012), The 2012 Pianura Padana Emiliana seismic sequence: Locations, moment tensors and magnitudes, Ann.Geophys., 55(4), 549–559, doi:10.4401/ag-6159.

Q8Segall, P. (2005) Earthquake and Volcano Deformation, John Wiley, Hoboken, NJ, USA.Shapiro, S. a., R. Patzig, E. Rothert, and J. Rindschwentner (2003), Triggering of seismicity by pore-pressure perturbations:

Permeability-related signatures of the phenomenon, Pure Appl. Geophys., 160(5), 1051–1066, doi:10.1007/PL00012560.Sitharam, T. G., J. Sridevi, and N. Shimizu (2001), Practical equivalent continuum characterization of jointed rock masses, Int. J. Rock Mech. Min.

Sci., 38(3), 437–448, doi:10.1016/S1365-1609(01)00010-7.Q9Sneed, M., D. L. Galloway, and W. L. Cunningham (2005), Earthquakes-rattling the Earth’s plumbing system, in Water Encyclopedia, John

Wiley, Hoboken, NJ, USA.Steacy, S. (2005), Introduction to special section: Stress transfer, earthquake triggering, and time-dependent seismic hazard, J. Geophys. Res.,

110, B05S01, doi:10.1029/2005JB003692.Stramondo, S., M. Saroli, C. Tolomei, M. Moro, F. Doumaz, A. Pesci, F. Loddo, P. Baldi, and E. Boschi (2007), Surface movements in Bologna (Po

plain—Italy) detected by multitemporal DInSAR, Remote Sens. Environ., 110(3), 304–316, doi:10.1016/j.rse.2007.02.023.Styles, P., P. Gasparini, E. Huenges, P. Scandone, S. Lasocki, and F. Terlizzese (2014), Report on the hydrocarbon exploration and seismicity in

Emilia region (ICHESE Report),, 213. [Available at http://mappegis.regione.emilia-romagna.it/gstatico/documenti/ICHESE/ICHESE_Report.pdf (Accessed 1 January 2016).]

Terzaghi, K. (1943), Theoretical soil mechanics, Géotechnique, 510, doi:10.1016/0167-1987(88)90005-0.Tizzani, P., et al. (2013), New insights into the 2012 Emilia (Italy) seismic sequence through advanced numerical modeling of ground

deformation InSAR measurements, Geophys. Res. Lett., 40, 1971–1977, doi:10.1002/grl.50290.Q10Wang, H. (2000) Theory of Linear Poroelasticity with Applications to Geomechanics and Hydrogeology, Princeton Univ. Press. Q11

Whitehead, R. L., R. W. Harper, and H. G. Sisco (1985), Hydrologic changes associated with the October 28, 1983, Idaho earthquake, Pure Appl.Geophys., 122(2–4), 280–293, doi:10.1007/BF00874599.

Journal of Geophysical Research: Solid Earth 10.1002/2017JB014009

ALBANO ET AL. THE 2012 EMILIA-ROMAGNA EARTHQUAKE 17

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758596061626364656667

Author Query Form

Journal: Journal of Geophysical Research: Solid Earth

Article: jgrb_52152

Dear Author,

During the copyediting of your paper, the following queries arose. Please respond to these by annotating your proofs with the necessarychanges/additions.• If you intend to annotate your proof electronically, please refer to the E-annotation guidelines.• If you intend to annotate your proof by means of hard-copy mark-up, please use the standard proofing marks. If manually writing cor-rections on your proof and returning it by fax, do not write too close to the edge of the paper. Please remember that illegible mark-upsmay delay publication.

Whether you opt for hard-copy or electronic annotation of your proofs, we recommend that you provide additional clarification of answersto queries by entering your answers on the query sheet, in addition to the text mark-up.

Query No. Query Remark

Q1 AUTHOR: Please complete the Publication Fee Consent Form included with your article andreturn to the Production Editor with your proofs.

Q2 AUTHOR: Please confirm that given names (red) and surnames/family names (green) have beenidentified correctly.

Q3 AUTHOR: Please check Table 1 if it is presented correctly.

Q4 AUTHOR: Please check figure captions if they are presented correctly.

Q5 AUTHOR: Please provide the page range for this chapter in Reference.

Q6 AUTHOR: Please provide the city location of publisher for Reference Fyfe, W. S. 2012.

Q7 AUTHOR: Please provide the page range for this chapter in Reference.

Q8 AUTHOR: Please provide the page range for this chapter in Reference.

Q9 AUTHOR: Please provide the page range for this chapter in Reference.

Q10 AUTHOR: Please provide the page range for this chapter in Reference.

Q11 AUTHOR: Please provide the city location of publisher for Reference Wang, H. 2000.

          

     

Additional reprint and journal issue purchases  

Should you wish to purchase additional copies of your article,  please click on the link and follow the instructions provided: 

9&acro= JCB Corresponding authors are invited to inform their co‐authors of the reprint options available. 

 Please note that regardless of the form in which they are acquired, reprints should not be resold, nor further disseminated in electronic form, nor deployed in part or in whole in any marketing, promotional or educational contexts without authorization from Wiley. Permissions requests should be directed to mailto: permissionsus@wiley.com 

 For information about ‘Pay‐Per‐View and Article Select’ click on the following link: http://wileyonlinelibrary.com/ppv  

 

https://caesar.sheridan.com/reprints/redir.php?pub=10089&acro=JGRB