Late Cenozoic uplift of western Turkey: Improved dating and numerical modelling of the Gediz river...

e 51 (2006) 131–171www.elsevier.com/locate/gloplacha

Global and Planetary Chang

Late Cenozoic uplift of western Turkey: Improved datingof the Kula Quaternary volcanic field and numerical

modelling of the Gediz River terrace staircase

Rob Westaway a,⁎, Hervé Guillou b, Sema Yurtmen c,1, Anthony Beck d,2,David Bridgland d, Tuncer Demir e, Stéphane Scaillet b, George Rowbotham f

a Faculty of Mathematics and Computing, The Open University, Eldon House, Gosforth, Newcastle-upon-Tyne NE3 3PW, UKb Laboratoire des Sciences du Climat et de l'Environnement, Domaine du CNRS,Bâtiment 12, Avenue de la Terrasse, 91198 Gif-sur-Yvette, France

c Department of Geology, Çukurova University, 01330 Adana, Turkeyd Department of Geography, Durham University, South Road, Durham DH1 3LE, UK

e Department of Geography, Harran University, 63300 Şanlıurfa, Turkeyf School of Earth Sciences and Geography, Keele University, Keele, Staffordshire ST5 5BG, UK

Available online 20 March 2006

Abstract

A set of 13 new unspiked K–Ar dates has been obtained for the Quaternary basaltic volcanism in the Kula area of westernTurkey, providing improved age control for the fluvial deposits of the Gediz River that underlie these basalt flows. This dating isable, for the first time, to resolve different ages for the oldest basalts, assigned to category β2, that cap the earliest Gediz depositsrecognised in this area, at altitudes of ∼140 to ∼210 m above present river level. In particular, the β2 basalt capping the SarnıçPlateau is dated to 1215±16 ka (±2σ), suggesting that the youngest underlying fluvial deposits, ∼185 m above present river level,are no younger than marine oxygen isotope stage (MIS) 38. In contrast, the β2 basalt capping the adjacent Burgaz Plateau is datedto 1014±23 ka, suggesting that the youngest underlying fluvial deposits, ∼140 m above present river level, date from MIS 28. Thestaircase of 11 high Gediz terraces capping the latter plateau is thus dated to MIS 48-28, assuming they represent consecutive∼40 ka Milankovitch cycles, although it is possible that as many as two cycles are missing from this sequence such that the highestterrace is correspondingly older. Basalt flows assigned to the β3 category, capping Gediz terraces ∼35 and ∼25 m above thepresent river level, have been dated to 236±6 ka and 180±5 ka, indicating incision rates of ∼0.15 mm a−1, similar to the time-averaged rates since the eruptions of the β2 basalts. The youngest basalts, assigned to category β4, are Late Holocene; our K–Arresults for them range from zero age to a maximum of 7±2 ka.

This fluvial incision is interpreted using numerical modelling as a consequence of uplift caused by a regional-scale increase inspatial average erosion rates to ∼0.1 mm a−1, starting at ∼3100 ka, caused by climate deterioration, since when a total of ∼410 mof uplift has occurred. Parameters deduced on this basis from the observed disposition of the Early Pleistocene Gediz terracesinclude the local effective viscosity of the lower crust, which is ∼2×1018 Pa s, the Moho temperature of ∼660 °C, and the depth ofthe base of the brittle upper crust, which is ∼13 km. The thin lithosphere in this area results in high heat flow, causing this relativelyshallow base of the brittle upper crust and the associated relatively thick lower-crustal layer, situated between depths of ∼13 and

⁎ Corresponding author. Also at: School of Civil Engineering and Geosciences, University of Newcastle-upon-Tyne, Newcastle-upon-Tyne NE17RU, UK.

E-mail address: [email protected] (R. Westaway).1 Present address: 41 Kingsway East, Westlands, Newcastle-under-Lyme, Staffordshire ST5 5PY, UK.2 Present address: Department of Archaeology, Durham University,South Road, Durham DH1 3LE, UK.

0921-8181/$ - see front matter © 2006 Elsevier B.V. All rights reserved.doi:10.1016/j.gloplacha.2006.02.001

mailto:[email protected]

http://dx.doi.org/10.1016/j.gloplacha.2006.02.001

132 R. Westaway et al. / Global and Planetary Change 51 (2006) 131–171

∼30 km. It estimated that around 900 ka, at the start of the ∼100 ka Milankovitch forcing, the spatial average erosion rate increasedslightly, to ∼0.12 mm a−1; the associated relatively sluggish variations in uplift rates are as expected given the relatively thicklower-crustal layer.

This modelling indicates that the growth of topography since the Pliocene in this study region has not involved a steady state.The landscape was significantly perturbed by the Middle Pliocene increase in erosion rates, and has subsequently adjusted towards—but not reached—a new steady state consistent with these increased erosion rates. It would not be possible to constrain what hasbeen occurring from the Middle to Late Pleistocene or even the Early Pleistocene uplift response alone; information regarding thestarting conditions is also essential, this being available in this region from the older geological record of stacked fluvial andlacustrine deposition. This result has major implications for the rigorous modelling of uplift histories in regions of rapid erosion,where preservation of information to constrain the starting conditions is unlikely.© 2006 Elsevier B.V. All rights reserved.

Keywords: Turkey; K–Ar dating; Quaternary; Uplift; Incision; Landscape evolution

1. Introduction

The Kula region of western Turkey is emerging as asuperb locality for studying fluvial incision and the as-sociated Quaternary environmental change in an uplift-

Fig. 1. Map of the upper reaches of the Gediz River in the vicinity of Uşak and(1997, Fig. 1). Insets show location. The right-hand inset also shows (as thick lof the Mesozoic Neotethys Ocean. Ophiolites associated with the more north

ing region (e.g., Westaway et al., 2003, 2004; Demiret al., 2004; Maddy et al., 2005). The Gediz, one of theprincipal rivers that drains westward into the AegeanSea (Fig. 1), has locally incised by ∼400 m into theformer land surface in this area, creating a valley system

Kula, based on Fig. 2 ofWestaway et al. (2004), adapted from Seyitoğluines with chevron ornament) Palaeogene (? Eocene) age sutures of armserly İzmir-Ankara suture are widespread in the present study region.

133R. Westaway et al. / Global and Planetary Change 51 (2006) 131–171

typically several kilometres wide, which is inset bymanyriver terraces. This incision has taken place in part intoareas of metamorphic basement, formed of the latestPrecambrian Menderes metamorphics (typically, gneiss,schist, quartzite and marble) and in part into Late Ceno-zoic fill of terrestrial sedimentary basins (notably, theSelendi and Uşak-Güre Basins; Fig. 1). In the southernSelendi Basin, around Kula, a sequence of basalt flowsfrom almost a hundred small volcanic necks (Fig. 2) haslocally backfilled this incised valley system, capping andthus preserving, or ‘fossilising’, the underlying fluvialdeposits (e.g., Ercan and Öztunalı, 1982; Richardson-Bunbury, 1996; Westaway et al., 2004; Maddy et al.,2005).

Following influential studies such as those by Jack-son et al. (1982) and Jackson and McKenzie, (1988), theview has been widely held that vertical crustal motions

Fig. 2. Map of the reach of the Gediz in the Kula area, adapted from Fig. 4 of WFig. 2) with additional information from Ozaner (1992) and Seyitoğlu (19Schist), marble, and quartzite, as well as chert and other lithologies fromdominated Hacıbekir Group forms the earliest part of the Late Cenozoic terreand is itself unconformably overlain by the basal part of the İnay Group (the BBasin; it is dated to the Early Miocene from its field relationships to dated volc2004, for more details). The inset shows the transverse profile A–A' (adapteillustrating typical field relationships between basalt flow units and river ter

within the Aegean extensional province relate only tothe active normal faulting that is occurring. Even re-cently, Bunbury et al. (2001) asserted that the incisionby the Gediz is in response to local footwall uplift, and isthus of local significance only. However, in the Kulaarea, the Gediz River is N20 km away from the nearestsignificant active normal fault zone (represented by theKırdamları Fault, bounding the northern margin of theAlaşehir Graben; Fig. 2). This distance exceeds thecharacteristic flexural wavelength of crustal deformationin the vicinity of normal faults in the Aegean region (cf.Westaway, 1993), so any fault-related vertical crustalmotions would be expected to die out well before theGediz is reached. Furthermore, although the margins ofthe Selendi and Uşak-Güre Basins were previouslythought to be bounded by Late Cenozoic normal faults(e.g., Seyitoğlu, 1997), the flanking uplands are now

estaway et al. (2004), based originally on Richardson-Bunbury (1996,97). ‘Metamorphic Basement’ shading includes schist (the Menderesthe ophiolite suite marking the İzmir-Ankara Suture. The sandstone-strial sequence in this region. It unconformably overlies the basementalçıklıdere Member of the Ahmetler Formation) in the central Selendianism elsewhere in the region (see Seyitoğlu, 1997, or Westaway et al.,d from Ozaner, 1992, Fig. 7) across the Gediz gorge near Kalınharmanraces.


regarded as relics of older crustal deformation thatbounded depressions in which terrestrial sediment lateraccumulated (e.g., Purvis and Robertson, 2004; West-away et al., 2004; Fig. 3a). Minor active normal faultinghas been reported by Purvis and Robertson (2004)within the Selendi Basin along the reach of the SelendiRiver between Selendi town and its Gediz confluence(Fig. 1), and Westaway et al. (2004) considered thepossibility that the atypically small net amount of inci-sion by the Gediz, since stacked deposition ceased, inthe vicinity of Palankaya at the western margin of thisbasin (Fig. 2) may also relate to minor local activenormal faulting. However, any such faulting seems lessimportant overall than other factors, such as the isostaticresponse to lateral variations in rates and amounts oferosion, in determining lateral variations in fluvial in-cision (Westaway et al., 2004). The principal variation isbetween the relatively broad valley systems that havedeveloped within the Selendi and Uşak-Güre Basins andthe much narrower gorges characteristic of incision intothe metamorphic basement. However, the pattern iscomplicated further because in some localities, such asdirectly north of Kula, the Gediz has incised through theentire sequence in the Selendi Basin and is now cuttinginto the underlying basement (locally, ophiolite relatedto the İzmir-Ankara suture of the former NorthernNeotethys Ocean; Figs. 1 and 2, inset). Nonetheless, itnow seems clear that this reach of the Gediz providesevidence that can constrain the history of regional upliftof western Turkey, not local complexities of normalfaulting that would make it of parochial significanceonly (e.g., Westaway et al., 2004).

As conventionally defined (after Ercan et al., 1978,1983), the stacked and typically subhorizontally beddedstratigraphy of the Selendi and Uşak-Güre Basins in-cludes the mainly clastic Ahmetler Formation overlainby the Ulubey Formation, which consists predominantlyof lacustrine limestone, these units collectively formingthe İnay Group. The Ahmetler Formation consists main-ly of the Balçıklıdere Member, comprising fluvial sandand typically ∼200 m thick, overlain by the silty and

Fig. 3. Field photos in the vicinity of Palankaya. (a) View NNW, from this rwhich is incising along the contact, at the western margin of the Selendi Basin,Member fluvial sand (to the right). It is evident from this exposure (and manormal fault plane (cf. Seyitoğlu, 1997). (b) View SSE across the Gediz gorgbasalt flow, ∼40 m above present river level of ∼275 m a.s.l., overlying Melinking Palankaya village to Palankaya bridge, the viewpoint for (a), which is35–40 m above the river. (c) Close-up view of the lower, platy-jointed, part oscale, at the site where basalt sample 01TR92 was collected. A few metres becan be observed beneath this basalt flow. (d) View NW, downstream alocharacteristic coarse polygenetic Gediz gravel forming its low (∼5 m) terrac

tuffaceous Gedikler Member, which can reach athickness of many tens of metres. Palaeocurrent datafrom the Balçıklıdere Member indicates palaeoflowfrom the south, revealing a palaeo-drainage-geometryunrelated to the modern west-flowing Gediz (Purvis andRobertson, 2004) and indicating that this ancestral riversystem transported material derived from erosion of theMenderesMassif farther south (Fig. 1). The ages of thesesediments have been the subject of major disputes (cf.Seyitoğlu, 1997; Westaway et al., 2004). In the Uşak-Güre Basin the uppermost Balçıklıdere Member issecurely dated to ∼7 Ma from its mammalian biostra-tigraphy and magnetostratigraphy (see summary of evi-dence by Westaway et al., 2004, who also listed originalsources) and we favour an equivalent age for thesedeposits in the Selendi Basin, where they are not directlydated, because the stratigraphy is so similar. In contrast,Purvis and Robertson (2004) and Purvis et al. (2005)have suggested an Early Miocene or early Middle Mio-cene age for the overlying Gedikler Member in theSelendi Basin from Ar–Ar dating of biotite and sanidinegrains from its tuffs. However, like Westaway et al.(2004) we consider it possible that these data indicateapparent ages much older than the true eruption age as aresult of inherited argon, this being accepted as a majorproblem with dating Late Cenozoic tuffs in other regions(e.g., Gansecki et al., 1998; Campbell et al., 2001);alternatively, the mineral grains may be detrital from thisregion's Early–Middle Miocene volcanism. A possiblereason for this switch at ∼7 Ma from sand to silt de-position in the Selendi and Uşak-Güre Basins is thatwhen extension began in the Alaşehir graben (at ∼7 Maaccording to Westaway et al., 2005) it disrupted theancestral northward river system that deposited theBalçıklıdere Member, causing the material subsequentlyeroded and transported northward from the MenderesMassif (Fig. 1) to be trapped within this graben. TheUlubey Formation is not directly dated. In principle, itshould ultimately be possible to tune its sedimentarysequence to the Milankovitch forcing of climate, as hasalready been done for other carbonate-rich Late

oad bridge at [PC 3745 8275], up the gorge of the Kızıldam tributary,between the Menderes schist basement (to the left) and the Balçıklıdereny others in the region) that this contact is an unconformity and not ae from the locality depicted in (d), showing the base of the Palankayanderes Schist basement. In the middle of the field of view is the tracklocally at the level of the base of the basalt, at∼310–315 m a.s.l., somef the basalt section above the track depicted in (c), with TD providinghind where he is standing, behind the exposed basalt face, fluvial sandng the Gediz from [PC 3900 8205] showing a section through thee.

Fig. 4. (a) Panoramic view of the eastern part of the Kula volcanic field, from a water-colour-tinted engraving in Fig. 11 of Plate 3 of Hamilton and Strickland (1841). (b) The same image, re-annotatedusing modern placenames and nomenclature. The viewpoint is in the vicinity of [PC 514 651] near Özdamları (⁎ in Fig. 2), just inside the metamorphic basement at the SE margin of the Selendi Basin,∼2 km north of necks 71 and 72. The downstream limit of the basalt flow unit from these necks, locally capping what is now known as the Balçıklıdere Member fluvial sand, is visible in theforeground.

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Cenozoic lake basins in the Aegean region (e.g., VanVugt et al., 1998, 2001; Steenbrink et al., 1999, 2000). Inthe absence of age control, taking account of the regionalcontext Westaway et al. (2004) inferred that its depo-sition ended around ∼3 Ma, being followed by a switchto fluvial incision accompanying a regional increase inuplift rates. In the Kula area itself, no evidence is knownto directly constrain the age of this switch from stackeddeposition to erosion and incision, although pockets ofthe Ulubey Formation limestone are found up to∼400 mabove present river level (see Westaway et al., 2004, fordetails, also below). However, farther upstream in theGediz this switch is revealed by a staircase of highterraces formed of cemented polygenetic Gediz gravel atup to ∼360 m above present river level, inset into theuppermost part of the underlying stacked sequence(see Westaway et al., 2004, for details, these depositsbeing illustrated in their Figs. 5, 8 and 13). The insetdisposition of these deposits was first recognised byWestaway et al. (2003, 2004); they were previouslymisinterpreted as the uppermost part of the stackedsequence of the Uşak-Güre Basin, known as the‘Asartepe Formation’ (cf. Seyitoğlu, 1997).

Hamilton and Strickland (1841) first recognised thatthe Kula basalts can conveniently be classified intothree categories by relative age (deduced from weath-ering) and altitude above the present level of the Gediz(Fig. 4a). The modern nomenclature, following Canetand Jaoul (1946) as modified by Richardson-Bunbury(1996), designates these categories as β2, β3 and β4(Fig. 4b). The youngest basalts, called β4, appear veryfresh in the field, their original exposed rubbly surfacesshow no subsequent soil development. These flowscascade down tributary gorges to the present level of theGediz, or to within a few metres of it; on each of thesegrounds they are evidently not very old. These freshbasalts are observed in three main localities (Fig. 2): inthe central part of the volcanic field around Gökçeören,where they do not reach the Gediz and so have not beenstudied by us; farther east where a flow from Divlit Tepe(neck 65; Fig. 4b) near Kula town reaches the Gediz atKula Bridge (Fig. 2); and, in the west, where anotherflow reaches the Gediz just downstream of DemirköprüDam. Below this dam this β4 flow follows the narrowGediz gorge in basement downstream for ∼6 km toAdala, crossing the Kırdamları Fault and entering theAlaşehir Graben (Fig. 2). This reach of the Gediz wasinvestigated by Westaway et al. (2004), being illustratedin their Figs. 11,19 and 20. Below Kula Bridge, the β4basalt continues down the Gediz gorge for almost 3 km,where as many previous studies (e.g., Hamilton andStrickland, 1841; Richardson-Bunbury, 1996) have

noted, it is inset into multiple generations of olderbasalt flows; this reach is illustrated in Figs. 12 and 17 ofWestaway et al. (2004) (see also Fig. 5).

The β3 category designates basalt that appears lessfresh in the field, being typically overlain by asufficiently well-developed soil to permit cultivation,but which also cascades down tributary valleys into themodern Gediz gorge, its base being typically a few tensof metres above present river level. As is illustrated inFig. 2, this β3 category represents the bulk of the Kulabasalt. It abuts the Gediz in four main localities: in thewest along and downstream of the Demirköprü res-ervoir, and progressively farther upstream around Palan-kaya (Fig. 3b,c), Kula Bridge (Figs. 4b and 6a) andDereköy (Fig. 6b). Some of the β3 category sites aroundDemirköprü reservoir were briefly described by West-away et al. (2004), but are not considered in the presentstudy; the other three localities are all investigated here.

In contrast with the younger basalt flows, those of theeven more weathered and dissected β2 category do notcascade down tributary valleys, and are not found within∼140 m of present river level (Figs. 4b, 6c). This β2basalt has been described in the past as ‘plateau basalt’,and is indeed found covering the İbrahimağa Plateaudownstream of Kula Bridge, in the left bank of the Gediz,and the Sarnıç and Burgaz plateaus farther upstream inits right bank (Fig. 2). It has long been recognised thatthese basalts cap thin gravels that overlie the BalçıklıdereMember sand (e.g., Ercan and Öztunalı, 1982; Richard-son-Bunbury, 1996). Westaway et al. (2003, 2004) notedthat much of this gravel is fluvial, and consists in part ofpolygenetic Gediz gravel and in part of limestone gravelderived from erosion of the Ulubey Formation by localtributaries. Westaway et al. (2004) also tentatively re-cognised fluvial terrace bluffs beneath this basalt, andinferred that the gravels in this vicinity record progres-sive lateral migration of the Gediz towards its presentcourse (southward across the Sarnıç and Burgaz plateausand northward across the İbrahimağa Plateau) during anearly stage of its incision history. This conclusion wasconfirmed by Maddy et al. (2005), who recorded astaircase of up to 11 Gediz terraces beneath these pla-teaus (Fig. 7) and developed more detailed palaeo-envi-ronmental reconstructions, regarding the limestonegravels as deposited by low-angle fans shed from re-sidual uplands of Ulubey Formation limestone (Fig. 2),after incision into this sediment had begun. The β2basalts thus did not erupt onto flat land surfaces, theyinstead cascade across these staircases of older terraces,providing the means of dating them, just as capping bythe younger basalts can date the lower terraces of theGediz within its modern gorge.


2. K–Ar dating

2.1. Background

Absolute dating of the Kula volcanism has proved tobe a challenging task. Westaway et al. (2004) discussedthe historical dataset from this region at length. Theydiscounted the limited luminescence dating by Göksu(1978) and considered the very old ages (up to ∼8 Ma)produced by Ercan et al. (1985) from whole-rock datingof basalt samples to reflect inherited radiogenic argon inflows that erupted much later. The only ‘historical’numerical age that they considered reliable was the1.1 Ma (with no stated error margin) K–Ar date fromBorsi et al. (1972) for a sample from an unknownlocality in one of the β2 basalt plateaus flanking theGediz. Sanver (1968) attempted to date the Kula volca-nism from its geomagnetic polarity. However, onlyflows in the β3 and β4 categories were sampled, allbeing normally magnetised and so evidently attributableto the Brunhes chron. In contrast, the Borsi et al. (1972)date and our own dating (described below) indicate thatthis technique could usefully be applied to the β2 flows,which straddle the Matuyama–Cobb Mountain–Matuyama–Jaramillo geomagnetic reversal sequence(cf. Tauxe et al., 2004).

The recent review, by Tokçaer et al. (2005), of this‘historical’ dating evidence reported three other whole-rock K–Ar dates that Westaway et al. (2004) did notconsider. The first, 0.3±0.4 Ma (±2σ), was for the β2basalt from neck 76 (or Delihasan Tepe; V in Fig. 2),∼200 m above the Gediz. The other two dates, by Ercanet al. (1997), were: 0.6±0.2 Ma (±2σ) for the β3 basalt,sampled from one of the scoria cones near Sandal (Q inFig. 2), which cascades northward into the Gediz valleyat Palankaya (N in Fig. 2); and 0.1–0.2 Ma for a sample,collected near Kula Bridge (I in Fig. 2), of the large flowunit of β4 basalt that likewise cascades northward to thislocality. The first of these dates is much younger thanany other date for β2 basalt in the region; the remaining

Fig. 5. Illustrations in the vicinity of Kula Bridge. (a) View across the Gediz gyielded our basalt sample 00YM19 in its left bank. Theβ4 basalt flow locally hbeing marked by∼30 cm of fluvial sand∼8m above the Gediz, as illustrated cthe Gediz gorge from a locality just north of the viewpoint in (a). The β4 bascoarse rounded cobbles of basalt and basement lithologies. About ∼1 km awaflowed up the Geren tributary gorge for∼500 m. The Gediz River remains conthe lower part of the Geren valley as illustrated in Fig. 2. (c) Section through cjust upstreamof the tip of theβ4 basalt flow, but downstream of the tip of theβ3been exposed as a result of incision by the lower reach of the Hayırlı, which fabe traced on Fig. 8) and locally follows the left flank of the β4 basalt flow (noGediz just downstream of its tip. The gravel is locally∼2 m thick (RW providWe presume that this gravel was most likely deposited during MIS 2; it evide

two are much older than other age determinations avail-able to Westaway et al. (2004) for the same flow units.As a result, none of these three dates can now be con-sidered reliable.

The modern dating literature on this region wasbegun by Richardson-Bunbury (1996), which reported 540Ar / 39Ar dates for 12 sample splits consisting of am-phibole phenocrysts. It is not clear why this particulardating approach was chosen, as amphibole is well-known to retain radiogenic argon, having a particularlyhigh closure temperature for this element (e.g., McDou-gall and Harrison, 1999), and thus to give numericalages that exceed the true eruption ages. Westaway et al.(2004) indeed considered that no more than 4 of these 12sample splits yielded meaningful ages, resulting in atotal of 3 40Ar / 39Ar dates. They also presented 6 newdates for 11 sample splits, determined at the ScottishUniversities' Environmental Research Centre (SUERC)laboratory at East Kilbride, Scotland, by applying theunspiked (or ‘Cassignol’) variant of the K–Ar techniqueto groundmass from which the phenocrysts had beenremoved. This laborious preparation method offers agreater likelihood of producing meaningful ages, al-though there is always a chance that some small pheno-crysts will remain in a sample, resulting in a numericalage that overestimates the true eruption age. Of these 6samples, 5 were considered by Westaway et al. (2004)to give meaningful ages, although one (for sample00YM30) was subject to considerable uncertainty dueto discordance between its two sample splits. Thus, atthe start of the present study we considered there to be atotal of 9 meaningful dates for this Kula volcanismfrom the K–Ar system: 1 from Borsi et al. (1972), 3from Richardson-Bunbury (1996), and 5 from West-away et al. (2004).

The main problem in K–Ar dating concerns mea-suring the 40Ar / 36Ar ratio, which is very large, anddiffers only slightly, often by less than 1%, between thesample (the unknown) and the standard (the present-dayatmosphere). In the conventional method of isotopic

orge from the road in its right bank at [PC 49344 77055] to the site thatas a rubbly top and a ropey base; it caps ophiolitic basement, the contactlose-up in Fig. 17b ofWestaway et al. (2004). (b) View northward downalt flow descends to river level as the gorge widens, its course lined byy (at s in Fig. 8) the basalt flow splits; the branch visible in the distancefined on the right flank of this basalt flow, its present course looping upoarse Gediz gravel overlying ophiolitic basement at [PC 49154 78725],flow (at q in Fig. 8) intowhich the younger flow is inset. This gravel has

rther upstream follows the left flank of the β3 basalt flow (its course cant visible in this photo as it is directly behind the viewpoint), joining thees scale), its top estimated as∼5 m above the present level of the Gediz.ntly pre-dates the β4 basalt flow that is inset into it.

http://dx.doi.org/doi:10.1029/2003GC000661

Fig. 6. Photos of sites on the Gediz farther upstream. (a) View NW along the Gediz from the vicinity of [PC 505 756], NE of Kalınharman. Basalt,categorised as β3, from one of the necks NE of Kula town is locally inset into basement, its top and base at ∼430 and ∼415 m a.s.l.,∼60 and ∼45 mabove the present river level. This site corresponds with the section line A–A' in Fig. 2. It marks the start of a∼2.5 km long reach of the Gediz, whichends at Değirmenler (Fig. 5b), where its gorge is constricted as it flows through ophiolite beneath the Miocene sequence of the Selendi Basin. (b)View of the west flank of the Söğüt gorge at Dereköy, showing β3 basalt overlying polygenetic Gediz gravel, adjacent to the site where basalt sample01TR101 was collected, which yielded a date of 180 ka. TD and Darrel Maddy provide scale. (c) View NNW from near the present∼375 m a.s.l. levelof the Gediz in the vicinity of [PC 518 748], showing the southern margin of the Sarnıç Plateau south of Çakırca and its cap of β2 basalt that we havedated to ∼1215 ka. The Gediz has locally insised ∼185 m since the eruption of this basalt. To the left of the field of view is a flat land surface at∼475 m a.s.l., which we presume to mark a former level of the Gediz. Our dating and uplift modelling (see main text) support an uplift rate in this areaof ∼0.15 mm a−1, suggesting that the Gediz probably occupied this level around MIS 16 (∼640 ka).


Fig. 7. Map (a) and cross-section (b) through the Gediz high terrace staircase, adapted from Figs 2 and 3a of Maddy et al. (2005). Subsequentfieldwork by Maddy et al. (in press) has increased the number of documented sites in Gediz gravel assigned to this staircase from 47 to 65, but onlyone of these extra sites, site 51 (discussed in the text), is shown; the others are all located in the Burgaz Plateau, and confirm the transverse profile ofterraces beneath it as illustrated in (b). Dashed lines in (a) show the positions of Gediz terrace bluffs as inferred by Maddy et al. (2005). Some of thesebluffs are directly observed in the field; others are interpreted from differences in altitude of widely separated gravel outcrops. Dotted lines show theprincipal localities at which interpretations of terrace bluffs were adjusted between the Maddy et al. (2005) and Maddy et al. (in press) schemes.Further major adjustments to this terrace scheme are required in the light of the new evidence reported in the present study.


dilution, this ratio is not measured directly for the sam-ple. Instead, argon from the sample is mixed with aknown amount of a ‘spike’ of ‘known’ isotopic con-centration, which is mainly 38Ar. The resulting 40Ar /38Ar and 38Ar / 36Ar ratios are both quite large, but smallenough to be measurable. The concentration of radio-genic 40Ar is then calculated from the amount of 38Ar

in the spike and the measured isotope ratios for themixture and the atmospheric argon (e.g., Dalrymple andLanphere, 1969, p. 251).

The principal difficulties in dating Quaternary basaltsby isotopic dilution concern the need for accurateknowledge of the isotopic composition of the spike andthe need to incorporate measurement errors to calculate

Table 1K–Ar dates for basalt samples On separate sheet

K (±1σ) (wt%) Split Weighted mean

Lab no. M(g)

40Ar⁎

(%)

40Ar⁎ (±1σ)(pmol g−1)

Age (±1σ)(ka)

40Ar⁎ (±1σ)(pmol g−1)

Age (±2σ)(ka)

β2; 01TR74; Toytepe-Armutboğazı; [PC 53605 77242]0.971±0.010 (1.0%) 6373 1.01981 6.554 2.068±0.014 (0.7%) 1228±15 (1.2%)

6398 0.93982 5.125 2.035±0.015 (0.7%) 1208±15 (1.2%) 2.053±0.010 (0.5%) 1219±27 (1.1%)

β2; 00YM22; Sarnıç-Dikilitaş; [PC 52687 76491]0.996±0.010 (1.0%) 6372 0.98671 7.991 2.095±0.013 (0.6%) 1212±14 (1.2%)

6397 0.97361 10.100 1.962±0.013 (0.7%) 1136±14 (1.2%) 2.029±0.009 (0.4%) 1174±26 (1.1%)

β2; 00YM25; Burgaz; [PC 57521 76229]1.287±0.013 (1.0%) 6534 0.98864 4.079 2.224±0.016 (0.7%) 996±12 (1.2%)

6551 1.01972 5.402 2.302±0.016 (0.7%) 1031±12 (1.2%) 2.264±0.011 (0.5%) 1014±23 (1.1%)

β2; 01TR97; I.brahimağa-Sarıyar Sırtı; [PC 45934 77361]

2.690±0.027 (1.0%) 6536 1.07235 23.035 4.684±0.025 (0.5%) 1004±11 (1.1%)6550 0.93901 22.161 4.637±0.025 (0.5%) 994±11 (1.1%) 4.660±0.018 (0.4%) 999±21 (1.1%)

β3; 01TR101; Dereköy, Söğüt left bank; [PC 53930 73484]2.640±0.026 (1.0%) 6424 0.98404 3.734 1.061±0.010 (0.9%) 232±3 (1.4%)

6456 1.02751 4.179 1.100±0.013 (1.1%) 240±4 (1.5%) 1.076±0.008 (0.7%) 236±6 (1.3%)

β3; 01TR95; Yusufağa ‘waterfall’; [PC 46130 75080]2.200±0.022 (1.0%) 6423 1.09917 2.768 0.668±0.008 (1.1%) 175±3 (1.5%)

6455 1.05038 4.459 0.703±0.008 (1.1%) 184±3 (1.5%) 0.686±0.005 (0.7%) 180±5 (1.3%)

β3; 01TR92; Palankaya-bridge track (NW, base); [PC 37227 82597]1.486±0.015 (1.0%) 6504 1.01507 3.802 0.454±0.007 (1.5%) 176±3 (1.8%)

6520 1.41739 4.675 0.469±0.007 (1.4%) 182±3 (1.7%) 0.462±0.005 (1.1%) 179±5 (1.5%)

β3; 01TR90; Palankaya-Çurukmehmet (NE, top); [PC 38659 81671]2.100±0.021 (1.0%) 6376 1.21791 2.188 0.629±0.009 (1.4%) 173±3 (1.7%)

6400 0.99926 2.402 0.662±0.011 (1.7%) 182±4 (2.0%) 0.641±0.007 (1.1%) 176±5 (1.5%)

β3; 01TR89; Palankaya-Yamantepe (SE, top); [PC 37653 81202]2.092±0.021 (1.0%) 6374 1.02977 1.258 0.608±0.008 (1.4%) 168±3 (1.8%)

6399 1.04516 1.237 0.628±0.010 (1.5%) 173±3 (1.8%) 0.615±0.007 (1.1%) 170±5 (1.5%)

β4; 00YM19; Değirmenler, Gediz left bank; [PC 4910 7770]3.055±0.031 (1.0%) 6472 0.94682 0.184 0.025±0.010 (42%) 5±2 (42%)

6482 1.27180 0.171 0.040±0.008 (20%) 7±1 (20%) 0.035±0.006 (18%) 7±2 (19%)

β4; 01TR87; Kızlarçeşmesi, tip of young flow; [PC 48884 79331]3.113±0.031 (1.0%) 6549 0.88929 0.082 0.021±0.011 (53%) 4±2 (53%)

6575 1.01289 −0.044 ND ZA

β4; 00YM13; Kula Bridge, Gediz left bank; [PC 48719 76509]2.972±0.030 (1.0%) 6535 1.03746 −0.027 ND ZA

6560 1.02407 −0.014 ND ZA

β4; 00YM11; Demirköprü; [PC 13687 74643]2.939±0.030 (1.0%) 6745 1.29857 −0.245 ND ZA

6760 1.50557 −0.096 ND ZA

Notes to Table 1:All dates are on groundmass. For each sample, the header indicates the category of basalt to which it is assigned, the sample number, the site name,and its UTM co-ordinates. Sample 00YM11 utilised spare groundmass separate left over from the dating by Westaway et al. (2004). Sample 01TR89came from the same place as sample 00YM30 of Westaway et al. (2004) but was prepared and analysed independently. All other samples are from



the uncertainty in the resulting age. Only a few percentof the 40Ar in a ∼1 Ma basalt sample with ∼1 wt.% K istypically radiogenic. The percentage error in an agedetermination depends on the percentage errors in theoriginal 40Ar / 38Ar and 38Ar / 36Ar ratios in the mixture,divided by the estimated percentage, R, of 40Ar that isradiogenic (e.g., Cox and Dalrymple, 1967; Dalrympleand Lanphere, 1969, p. 105). The error analysis by Coxand Dalrymple (1967) indicates that this margin oferror will increase dramatically for younger sampleswith even smaller values of R (see, e.g., Dalrymple andLanphere, 1969, p. 106). To accurately date Middle toLate Pleistocene volcanic rocks, one must thus reliablymeasure very small values of R, of b1%. Lanphere(2000) has demonstrated the ability of the ‘spiked’K–Ar method to date rocks of this age, but it requireslarge (∼25 g) molten samples and accurate correctionfor the 40Ar / 36Ar ratio in the atmospheric argon in thelaboratory.

An alternative K–Ar method, involving smaller mol-ten samples (∼1–3 g), has proved successful in datingMiddle and Late Pleistocene volcanic rocks (Gillot andCornette, 1986; Guillou et al., 1997, 1998). This un-spiked technique depends on accurate calibration ofthe response of the mass spectrometer throughout therange of measurement of both the rare isotope 36Arand the more abundant 40Ar. By dynamically compar-ing the isotopic composition of an aliquot of pure atmo-spheric Ar with the sample Ar, it accurately determinesminor variations of the 40Ar / 36Ar ratio between thestandard and the unknown. This method thus providesa precise correction for atmospheric argon contamina-tion and also avoids any discrimination effects of themass spectrometric measurement (Cassignol and Gillot,1982).

40Ar / 39Ar dating is, of course, another methodutilizing the K–Ar system, which has been more widelyapplied to very young rocks (e.g., McDougall andHarrison, 1999). Comparisons between unspiked K–Ar

Notes to Table 1:All dates are on groundmass. For each sample, the header indicates the categand its UTM co-ordinates. Sample 00YM11 utilised spare groundmass separacame from the same place as sample 00YM30 of Westaway et al. (2004) butnew sites that have not previously been dated. K indicates the concentration omolten mass of each sample split. 40Ar⁎ (%) indicates the calculated percentagmeasured argon isotope ratios) to be radiogenic. 40Ar⁎ (pmol/g) indicates thethis column indicates that the concentration could not be determined becausecalculated age for each sample split, treated individually. Weighted mean agethe text. ZA in either age column indicates that the sample is assigned a nomiAge calculations are based on the decay and isotopic abundance constantsmeasurements and in age determinations are expressed in both absolute and pindividual measurements combine to give the resulting percentage uncertainuncertainty in the overall age determination is calculated for ±1σ even thou

analyses and Ar–Ar analyses on similar Quaternarysamples (Guillou and Singer, 1997; Singer et al., 2004;Guillou et al., 2001, 2004) reveal them to be consistent,with similar error bounds. They thus indicate the po-tential of the unspiked method, justifying our use of it todate Quaternary basalts.

2.2. Laboratory procedure

Samples, typically weighing ∼500 g and with noalteration or only minimal alteration evident in handspecimen, were sawn to allow internal surfaces of anyvesicles to be inspected for evidence of alteration prod-ucts. Inspection of petrographic thin sections furtherrefined the sample selection. Samples were then crush-ed, washed in deionized water and dilute hydrochloricacid, sieved to a 60–80 μm size fraction, and pheno-crysts and xenocrysts were removed using magneticseparation and hand picking. Potassium and argon weremeasured in this microcrystalline groundmass. Thisprocess improves the potassium yield as well as thepercentage of radiogenic argon (Guillou et al., 1998)and removes the main potential source of systematicerror due to the tendency of excess, or ‘inherited’, ra-diogenic 40Ar to occur in phenocrysts.

Samples were K–Ar dated at the Laboratoire desSciences du Climat et de l’Environnement (LSCE) atGif-sur-Yvette, France (a joint division of the CentreNational de la Recherche Scientifique and the Commis-sariat à l’Energie Atomique), using the unspiked (orCassignol) method (e.g., Cassignol et al., 1978;Cassignol and Gillot, 1982; Gillot et al., 1982), whichis based on the fundamental assumptions of conven-tional K–Ar dating as defined by Dalrymple andLanphere (1969). Crystallisation ages may be under-or overestimated in some cases if rocks have not evolvedas a closed system for K and Ar since solidification. Thepurpose of our selection of unaltered rocks by petrog-raphic examination was to reduce this source of error.

ory of basalt to which it is assigned, the sample number, the site name,te left over from the dating by Westaway et al. (2004). Sample 01TR89was prepared and analysed independently. All other samples are fromf potassium in the sample expressed as weight percent of K2O. M is thee of the 40Ar in the sample that is determined (by calculation, using thecorresponding concentration of radiogenic argon in the sample. ND inthe calculated percentage of radiogenic argon was negative. Age is theis a best estimate of the age of each sample, calculated as explained innal ‘zero age’ because no radiogenic argon could be resolved within it.from Steiger and Jäger (1977). Uncertainties in potassium and argonercentage terms. The latter clarify how the percentage uncertainties inty in each age determination. To facilitate comparison, the percentagegh the absolute error is expressed as ±2σ.


Undegassed, magmatic argon is another potential sourceof inaccurate ages. As mentioned above, many studies(e.g., Lanphere and Dalrymple, 1976; Kaneoka et al.,1983; Laughlin et al., 1994; Singer et al., 2004) havedemonstrated that phenocrysts may be carriers of excessargon. The magnetic separation prior to dating mini-mises (but does not eliminate) the probability of erro-neously old ages due to this source of excess 40Ar.

The most critical uncertainty in the K–Ar method isthat it is not possible to verify the isotopic compositionof the initial argon in the samples. That is, we cannotcheck the assumption that, at its time of formation, the40Ar / 36Ar ratio in any sample was the modern atmo-spheric value (295.5). As a result, the analytical errorsgiven in Table 1 may in some cases be less than the realerror.

Potassium was analysed by atomic spectrophotom-etry with a relative precision of 1%. Argon was extract-ed by radio-frequency heating from ∼1 g splits of thegroundmass samples, introduced into a high-vacuumglass line, and purified with titanium sponge and SAESZr–Al getters. Isotopic analyses were performed on Arquantities ranging from ∼10 to ∼70 pmol. Theinstrumental atmospheric argon correction was checkedby repeated measurements of zero-age standard samples(i.e., atmospheric blanks), which indicate a lower limitof 40Ar⁎ detection (δi; see below) of 0.14% (Scaillet andGuillou, 2004). The experimental procedure isexplained in more detail by Charbit et al. (1998).

Our error analysis incorporates uncertainties in po-tassium content, K, mass spectrometer (MS) calibration,and radiogenic argon yield. We define f(K) as the frac-tional error in K in each sample, this parameter beingassigned a nominal value of 0.01% or 1% for all ana-lyses. We define f(C) as the fractional error in the MScalibration C (the scale factor for converting millivoltsof MS output to atoms of argon), with a nominal valueof 0.005 or 0.5%. We also define, δi, the yield ofradiogenic argon in the MS run for sample split i, as

di ¼ Rui−RaiRui

; ð1Þ

where Ru and Ra are the 40Ar / 36Ar ratios measured inthe sample and in the associated atmospheric blank,respectively. The fractional variance in δi is thus f

2(δi),

f 2ðdiÞ ¼ RaiRui−Rai

� �ðf 2ðRaiÞ þ f 2ðRuiÞÞ ð2Þ

where

f 2ðRuiÞ ¼ r2ðRuiÞ=Ru2i ð3Þ

and

f 2ðRaiÞ ¼ r2ðRaiÞ=Ra2i ; ð4Þ

σ2(Rui) and σ2(Rai) being the variances in Rui and Rai,measured during the MS runs for the ith sample split andthe associated atmospheric blank.

The radiogenic argon content from MS run i is givenby

40Ar⁎i ¼ 40ArT i � di � C ; ð5Þ

where 40Ar iT is the total yield of 40Ar for this run. Its

fractional variance is thus

f 2ð40Ar⁎iÞ ¼ f 2ð40ArT iÞ þ f 2ðdiÞ þ f 2ðCÞ; ð6Þ

assuming that the uncertainties in 40AriT, δi, and C are

uncorrelated. The overall content of radiogenic argon ineach sample, determined as the weighted mean of n MSruns, is thus

b40Ar⁎N ¼ 1w

Xni¼1

40 Ar⁎ir2ð40Ar⁎i Þ2

ð7Þ

where

r2ð40Ar⁎iÞ ¼ f 2ð40Ar⁎i Þ � ð40Ar⁎iÞ2 ð8Þis the variance of the ith measurement of radiogenic40Ar, and acts as the weighting factor for the corre-sponding sample split. The overall weighting factorW isthus

W ¼Xni¼1

1

r2ð40Ar⁎iÞ; ð9Þ

and the variance and fractional error in b40Ar⁎N are

r2ðb40Ar⁎NÞ ¼ 1=W ð10Þand

f ðb40Ar⁎NÞ ¼ rðb40Ar⁎NÞ=b40Ar⁎N: ð11Þ

Using standard theory, such as Eq. (4–2) ofDalrymple and Lanphere (1969), the best estimate ofthe sample age can thus be determined as:

t ¼ 1ke þ ke Vþ kb

ln 1þ ke þ ke Vkbke þ ke V

b40Ar⁎N40K

� �ð12Þ

where 40K is the concentration of 40K in the sample(calculated using K, above, and the isotopic abundancefrom Steiger and Jäger, 1977), λε and λε′ are the decay


constants for the two modes of decay of 40K to 40Ar byelectron capture, and λβ is the decay constant for the betadecay of 40K to 40Ca. The variance in this age deter-mination is thus:

r2ðtÞ ¼ 1þ kke þ ke V

b40Ar⁎N40K

� �−2

� 1ke þ ke V

b40Ar⁎N40K

� �2�f 2ðb40Ar⁎NÞþf 2ðKÞ

�

ð13Þ

where λ=λε+λε′+λβ.Sample ages and their uncertainties, calculated on

this basis, are listed in Table 1. The 13 new K–Ar datesobtained comprise 4 for β2 basalt, 5 for β3 basalt and 4for β4 basalt. Eleven of these dates are for new sites, one(01TR89) is from the same site as sample 00YM30 ofWestaway et al. (2004), which yielded discordant splits,and one (00YM11) involves dating of spare materialfrom the Westaway et al. (2004) study as an inter-lab-oratory control.

2.3. The β4 flow units

The only dating considered meaningful by Westawayet al. (2004) for any of the β4 basalt flows was for thethree samples they dated: two (00YM11 and 00YM12)from the Demirköprü–Adala area and one (00YM17)from near Kula Bridge (Fig. 2). All three dates were inthe range ∼50–90 ka, the ages of the six individualsample splits being concentrated around 60 ka.

For the present study, sample 00YM13 was collectedfrom the west side of the Kula-Selendi road where itdescends along a cutting through the β4 flow, directlysouth of Kula Bridge, as illustrated in Fig. 17a ofWestaway et al. (2004) (see Fig. 8). Sample 00YM19was collected a few hundred metres farther downstream(Fig. 8) near the small settlement of Değirmenler, alsofrom the left bank of the Gediz, where this basalt caps a∼30 cm thick bed of fluvial gravelly sand (illustrated inFig. 17b of Westaway et al., 2004; see also Fig. 5a),∼8 m above the present river level. Sample 01TR87came from the tip of this β4 flow unit, in the valley flooron the left side of the Gediz, ∼2 km farther downstream,adjacent to its confluence with the small Kızlarçeşmesitributary (Figs 5c and 8). We also re-dated spare materialof sample 00YM11 of Westaway et al. (2004), frombelow Demirköprü Dam (Fig. 2) (see their Figs. 11 and19a for more detailed location information).

In contrast to the dating by Westaway et al. (2004),dating of the eight splits from our four samples (Table 1)

yielded three with ages of a few thousand years and fiveof zero age (i.e., in which no radiogenic argon could beresolved, δi in Eq. (1) having yielded negative values).One possible explanation for this difference, relative tothe results of Westaway et al. (2004), might be the use ofdifferent laboratory procedures for correcting the atmo-spheric argon contamination. In the SUERC laboratorythat ran the Westaway et al. (2004) samples, the proce-dure at the time involved only intermittent measure-ments of zero-age air standards to check the atmosphericcomposition of the instrumental blanks for input intoEq. (1). In contrast, the LSCE laboratory runs alternateair standards and dated samples. Generally, thisprecaution results in much more accurate estimates ofthe isotopic composition of the atmospheric argon in thelaboratory, which varies significantly due to memoryeffects and to the non linear response of the massspectrometer. Given the available geological evidence,and in accordance with the K–Ar data presented here,we deduce that these β4 basalt flows are Holocene, botharound Kula Bridge and in the Demirköprü–Adala area(cf. Yılmaz, 1990). Field evidence also indicates thatthese β4 basalt flows are very young, because in almostall localities where they are observed they are at thepresent river level (see Fig. 19a,b of Westaway et al.,2004, also Fig 5b); localities where the base of this basaltis above present river level, such as in Fig. 5a, are clearlythe exception rather than the rule.

A young age bound for these β4 basalts is providedby the absence of any historical record of volcanismaround Kula; we consider that this precludes anyeruption in the past four millennia, for which writtenrecords exist. Kula is indeed not listed in catalogues ofactive volcanic fields such as Blumenthal et al. (1964).However, the discovery of fossil human footprints intephra from a β4-category eruption near DemirköprüDam (e.g., Ozansoy, 1972; Barnaby, 1975) indicateshuman occupation of the area at the time. It is thought(e.g., French, 1961, 1969; Yakar, 1991, pp. 176–179)that the Holocene human presence in western Turkeybegan in the Aceramic Neolithic era starting at ∼7000B.C. (∼9 ka); the evidence becomes more abundant inNeolithic (from ∼6000 B.C. or ∼8 ka onwards) andChalcolithic (from ∼5000 B.C. or ∼7 ka) times. How-ever, there is no other evidence for a human presence inthis reach of the Gediz and the flanking uplands duringthe early Holocene, early settlement of western Turkeybeing largely confined instead to alluvial plains thatwere suited to agriculture and provided natural migra-tion and trade routes, such as the interior of the AlaşehirGraben (e.g., French, 1961, 1969). The Kula area wasfirst interpreted as an extinct volcanic field by Strabo of


Amaseia (Amasya) who lived during 65 BC to AD 23, ithaving been regarded earlier in antiquity as one of theentrances to the underworld (e.g., Jones, 1954). Onemay speculate that this superstition may have resultedfrom early inhabitants witnessing the eruptions thatproduced the β4 basalt flows.


Three K–Ar and Ar–Ar dates for the Kula volcanismclassified as β3 were considered meaningful by West-away et al. (2004). First, Richardson-Bunbury (1996)dated to 130±180 ka (±2σ) a single sample split from aβ3 basalt flow from neck 59 (Fig. 2), which she re-garded as cascading into the Gediz gorge around KulaBridge (reaching locality q in Fig. 8). Second, she datedto 190±100 ka (±2σ) another single sample split from aβ3 basalt flow from neck 32 (Fig. 2), which sheregarded as cascading into the Gediz gorge at Palankaya(Figs. 3b,c, and 9). Third, Westaway et al. (2004) datedtheir sample 00YM30 from the top of this Palankayaflow unit (Fig. 2) to 205±26 ka (±2σ). Although con-cordant with Richardson-Bunbury's (1996) date for thesame flow unit, this age was based on two sample splitsthat individually yielded discordant ages of 145±32 and299±32 ka (±2σ).

This Palankaya flow unit (Figs 3,9) is significant fordetermining the rate of incision by the Gediz, being thehighest-level (relative to the Gediz) β3 flow unit that isknown to cap sand and gravel of this river and with aclear geometrical relationship to its palaeo-course at thetime of eruption (Westaway et al., 2004). This flow unitis∼60 m thick in the vicinity of the Gediz gorge, its baseand top at estimated levels of ∼40 and ∼100 m abovethe present river level (see Fig. 3b,c and its caption, alsoFig. 9). As noted by Westaway et al. (2004), it cascadeddown the valley of the Yamantepe left-bank tributary,spreading out both upstream and downstream along theGediz for a total distance of ∼2.5 km and temporarilydamming it at the point where its valley abruptly nar-rows on passing out of the Selendi Basin and intometamorphic basement (Figs. 3b and 9). The Gediz hassubsequently incised around the northern margin of this

Fig. 8. Representations of the topography in the vicinity of Kula Bridge. (a) L(b) Contour map (contour interval 5 m) derived from Shuttle Radar TopograWestaway et al., 2006, and references therein, for more details), illustratingdenote sites discussed by Westaway et al. (2004) or this study. Large two- ordated. Digits followed by letters indicate viewpoints for Figs 5a–c and otherdiscussed in the text or figure captions, and include for informaton some of thin the present manuscript. Labelling of spot heights and contours in black indgrey (where shown) are derived from HGK topographic maps and (as discussefor sources of imagery and processing techniques.

flow unit, except directly opposite Palankaya village(g in Fig. 9b) where a small outlier of basalt overliesbasement in the modern right bank of the Gediz.

Given the importance of this Palankaya flow unit, wehave dated three new samples from it. The first, 01TR89,from its top at its SE margin where it is well-exposed byincision by the Yamantepe (as illustrated in Fig. 18b ofWestaway et al., 2004; see also Fig. 9), was intended toreplicate Westaway et al. (2004) sample 00YM30. It wascollected as close as possible to the same place, andyielded a date of 170±5 ka (Table 1). Second, sample01TR90 was collected from the top of this flow unit atits NE margin, the Çurukmehmet cliff (illustrated inFig. 18a of Westaway et al., 2004), which has formed asa result of subsequent incision by the Gediz. This sampleyielded a date of 176±5 ka. Third, sample 01TR92 wascollected from the base of this flow unit where it is well-exposed along the track linking Palankaya village to thelocal Gediz bridge, opposite the small right-bank basaltoutlier (Figs 3b,c). The basalt is locally observed tooverlie fluvial sand and above its chilled base displaysplaty jointing, giving way higher up to impressivecolumnar jointing indicative of relatively slow cooling,suggesting that a single eruption produced this entire∼60 m thickness of basalt. Sample 01TR92 yielded adate of 179±5 ka, not significantly different from thesamples from the top of this flow unit, confirming that itindeed formed in a single eruption, with an estimatedage—taken as the mean of our three dates and excludingboth previous dates—of 175±3 ka (±2σ). A time-averaged incision rate of∼0.23mm a−1 is thus indicated.

We wanted to reinforce Richardson-Bunbury's(1996) date for the β3 basalt at Kula Bridge (Fig. 8)by sampling this flow unit as close as possible to theGediz. This is because her date came from a site at aneck, ∼7 km from the Gediz; as noted by Westaway etal. (2004) there are so many necks and β3 flows in thisarea (Fig. 2) that there can be no certainty that the flowunit that she dated is the same one that reaches theGediz. As also noted by Westaway et al. (2004), besidethe Gediz in the vicinity of Değirmenler (at site q inFig. 8) this β3 flow unit can be observed to cap Gedizgravel, its base being ∼25 m above the present river

andsat ETM+ satellite image, panchromatic band with 15 m resolution.phic Mission (SRTM) radar imagery (see, e.g., Rabus et al., 2003, orits 90 m resolution. In this and in subsequent Figs., letter codes in (b)three-digit numbers indicate locations of basalt samples that we have

field photographs. Single letters and single-digit numbers indicate sitese principal sites described by Westaway et al. (2004) but not mentionedicates altitudes derived from the SRTM data; corresponding values ind in the text) are typically∼10±5 m lower. See Westaway et al. (2006)

Fig. 9. Similar imagery to Fig. 8 but for the vicinity of Palankaya.



level. However, the β3 basalt is locally highly altered,presumably through contact with river water, making itnot worth dating. We instead sampled this flow unit∼3 km from the Gediz (site o in Fig. 10), near Yusufağafarm, adjacent to the junction of the main Kula-Selendiroad with the minor roads to İbrahimağa and Ahmetli.As noted by Westaway et al. (2004), in this locality oneof several quite abrupt ‘steps’ in the surface of the β3basalt can be observed (notably where the Selendi roaddescends it, at site î in Fig. 10), where it descends by∼40 m in ∼300 m distance—possibly because thisbasalt has cascaded down a river terrace bluff or thefront of an older basalt flow that is now buried. At site othe Karakam tributary of the Gediz has locally incisedinto the β3 basalt, forming a seasonal waterfall and ex-posing a fresh section to permit sampling. Sample01TR95 from this site yielded a date of 180±5 ka, suchthat the observed ∼25 m of incision of this flow unit bythe Gediz indicates a time-averaged incision rate of∼0.14 mm a−1. This new date is concordant with thatpreviously obtained by Richardson-Bunbury (1996) butmuch more tightly constrained. Furthermore, as thereare no necks and no known flow fronts between this siteand the Gediz one can be confident that our new date isappropriate for the β3 basalt in the Gediz gorge itself.

Our final site in β3 basalt is at the eastern end of theKula volcanic field near Dereköy (Figs. 2, 6b, and 11)).Basalt from neck 68 has locally flowed NNE for ∼7 kmdown the Söğüt tributary valley to the Gediz, where itspread laterally both upstream and downstream to a totaldistance of ∼4 km, similar to the observed situation atPalankaya. The Gediz has subsequently re-established itscourse around the northern limit of this basalt. For muchof its length the Söğüt has reincised along the easternmargin of this flow unit, somewhat like the Yamantepe atPalankaya (Fig. 9). However, near the present course ofthe Gediz the Söğüt has instead incised through the basaltand into the underlyingBalçıklıdereMember fluvial sand,also exposing Gediz gravel, as illustrated schematicallyfor an adjacent locality in cross-section A–A′ in Fig. 2.

This succession is particularly clear ∼1 km from themodern course of the Gediz in the left bank of the Söğütgorge (Fig. 11), where sample 01TR101 was collected(Fig. 6b). This gorge is locally ∼40 m deep, exposingthe uppermost ∼15 m of the Balçıklıdere Member sand,then∼0.5 m of fluvial gravel and∼25 m of basalt with a∼0.5 m thick chilled lower margin. The valley floor islocally ∼400 m a.s.l., ∼20 m above the level of theGediz, ∼1 km to the north. The exposed fluvial gravel iswell-sorted and fines upwards, with clasts of basalt,schist, quartzite, orange-red chert and marble (Fig. 6b).Its polygenetic character makes it clear that it was de-

posited by the Gediz, not a local tributary, which is thusestimated to have incised locally by ∼35 m since thedeposition of this gravel. Gediz gravel at a similar levelis also observed a short distance upstream beneath thebasalt forming the right flank of the Söğüt gorge. Fartherupstream along the Söğüt, this gravel is absent; thebasalt instead rests directly on the Balçıklıdere Mem-ber sand, as shown schematically in Fig. 2. Sample01TR101, from the base of this flow unit, yielded a dateof 236±6 ka (Table 1), indicating a time-averagedincision rate of ∼0.15 mm a−1.


Prior to the present study, three dates for the β2volcanism were considered meaningful. First, Borsi etal. (1972) reported an 1100 ka K–Ar date, but with noerror margin or site co-ordinates. Second, Richardson-Bunbury (1996) determined Ar–Ar dates of 1370±100and 1120±60 ka for two splits of a sample from neck 75at Burgaz (Fig. 2), which yield a weighted mean age of1186±78 ka. Third, Westaway et al. (2004) datedsample 00YM23 from Çakırca, in the western part of theSarnıç Plateau (Figs 2,6c), to 1264±15 ka. The basalt atthis site overlies polygenetic Gediz gravel (illustrated inFig. 14e of Westaway et al., 2004) that has been as-signed to Gediz high terrace I of Maddy et al. (2005).

Our first new K–Ar date for this β2 volcanism is forsample 00YM22, collected from a small exposure be-side the road across the Sarnıç Plateau from Sarnıç toÇakırca in the area known as Dikilitaş. On the westernedge of Sarnıç (site 21 in Fig. 2) a quarry beside thisroad exposes limestone gravel at ∼570 m a.s.l. (withclasts derived from the Ulubey Formation, indicatingdeposition by a local right-bank tributary of the Gediz;see Fig. 14d of Westaway et al., 2004) overlain bypalagonitic tuff (derived from the adjacent neck 74 orSarnıç Bağtepe, indicating rapid cooling and thussuggesting that the site was at or near the contempora-neous river level). Maddy et al. (in press) found anadjacent site (site 51, assigned to their high terrace VI;Fig. 7a) where Gediz gravel is visible below thislimestone gravel, at an altitude of ∼565 m a.s.l.. Gedizgravel has also been found ∼400 m farther WNWat site26 of Maddy et al. (2005; Fig. 7a), this deposit being at∼560 m a.s.l. and assigned to their high terrace V.Subsequent eruptions of basalt from this neck raised thelevel of the land surface by up to∼70–80 m above thesedeposits (its summit is now at 637 m a.s.l.). Sample00YM22 was collected from ∼1 km west of site 21,close to the col in the middle of the Sarnıç Plateau, wherethe pattern of gentle topographic gradients suggests that


basalt flowing NW or WNW from Sarnıç Bağtepecoalesced with that flowing southward or SW fromToytepe (neck 73) at the northern plateaumargin (Fig. 2).It is not obvious from which of these two necks sample00YM22 originated, but our dating assigns it an age of1174±26 ka (Table 1).

Second, sample 01TR74 was collected from thin (nomore than ∼4 m thick) basalt from Toytepe neck, at theeastern margin of the Sarnıç Plateau (Fig. 2) overlookinga highly dissected area of Balçıklıdere Member fluvialsand called Armutboğazı, overlying cemented Gedizgravel (illustrated in Fig. 14c of Westaway et al., 2004),with local limestone (fan) gravel in between. This site isadjacent to site 31 of Maddy et al. (2005) (Fig. 7a), forwhich the fluvial gravel is assigned to their Gediz highterrace VI. Toytepe neck, ∼800 m farther north, fromwhich this basalt originated, is illustrated in Fig 14a andb of Westaway et al. (2004). Sample 01TR74 has pro-vided a K–Ar date of 1219±27 ka.

Our third β2 sample, 00YM25, was collected at∼620 m a.s.l. just north of Burgaz village (Fig. 2); itevidently erupted from the adjacent neck 75, or BurgazBağtepe. Basalt from this neck evidently flowed south-ward for∼2 km farther, capping the whole surface of theBurgaz plateau as well as adjacent smaller basalt-cappedhills such as Kale Tepe (Figs. 2, 7a), which are nowseparated from the main plateau by fluvial incisionthrough this basalt and the underlying BalçıklıdereMem-ber sand (this badland landscape at Kale Tepe is illus-trated in Fig. 3 ofWestaway et al., 2004). BurgazBağtepeneck may well also have produced the basalt caps of theresidual flat-topped hills located between the Burgaz andSarnıç Plateaus, notably Kavtepe (X in Fig. 2, at∼600 ma.s.l.; see Fig. 9 ofWestaway et al., 2004, for more detail)and İnkale Tepe (Y in Fig. 11, at 584 m a.s.l.). Thegeomorphology also suggests the possibility that thebasalt forming the eastern part of the Sarnıç Plateau,known as Sürtmece (Fig. 11), may have originated fromBurgaz Bağtepe neck and not from the adjacent SarnıçBağtepe neck. The upper surface of this basalt in theseareas indeed appears to have a continuous westwardcomponent of slope, from591m a.s.l. at Alaca (in the SWcorner of the Burgaz Plateau; Fig. 11) to 584 m a.s.l. atİnkale Tepe and 577 m in the eastern part of Sürtmece(Fig. 11), as illustrated in Fig. 16a and b of Westaway etal. (2004).

Westaway et al. (2004) noted several localities wherethis Burgaz Bağtepe basalt caps Gediz gravel, and sub-

Fig. 10. Similar imagery to Fig. 8 but for the vicinity of İbrahimağa. 1–6 markheights from HGK 1:25,000 scale topographic maps followed by those froN735 m; 3, unnamed, 743 /N735 m; 4, Eğlence Tepe, 727 /N720 m; 5, Abaş

sequently Maddy et al. (2005, in press) have mappedand surveyed these and many other localities systemat-ically, revealing a Gediz high staircase of 11 terraces(Fig. 7). This survey indicated that the lower altitudelimit of this basalt is ∼540 m a.s.l., on Kale Tepe,capping high terrace I of Maddy et al. (2005), ∼140 mabove the present level of the Gediz. Before this surveywas carried out, Westaway et al. (2004) estimated thisaltitude limit higher, ∼560 m, in the southern part of theBurgaz Plateau, in what is now designated as highterrace IV; until the area was surveyed it was not clearthat the base of the basalt is up to ∼20 m lower in themore fragmentary exposures farther south. Our sample00YM25 yielded a date of 1014±23 ka, which evidentlypost-dates this entire high terrace staircase (Fig. 7b).

Our final dated sample, 01TR97, came from thesouthernmargin of the İbrahimağa Plateau (Fig 10), fromthe cutting on the road across this plateau linkingKavakalanı and İbrahimağa, as it begins to descend,∼500 m east of the latter village, in the locality known asSarıyar Sırtı. As noted byWestaway et al. (2004), most ofthe basalt cap of this plateau has erupted from neck 57, orİbrahimağa Bağtepe, from which it flowed eastward for∼3.5 km and northward for ∼2.5 km, infilling a ∼4 kmlength of the Gediz palaeo-valley (Figs. 2 and 10). Someof the basalt forming the NE part of this plateau eruptedfrom the smaller neck 58, or Tavşan Tepe (Figs. 2 and10). Westaway et al. (2004) inferred that this plateau isunderlain by a succession of Gediz palaeo-courses,which record northward channel migration, in contrastwith the southward migration evident from the fluvialdeposits underlying the Burgaz and Sarnıç Plateaus.However, no fluvial deposits relating to any such palaeo-courses have yet been identified beneath the İbrahimağaPlateau. This is because investigation of its margins ismuch more difficult than for the Burgaz Plateau, asvegetation cover is more pervasive and these marginshave weathered to relatively gentle slopes that arecovered with basalt talus, making it difficult to examinethe contact between the base of the in situ basalt and theunderlying sediment. In the east near Burşuk, theİbrahimağa plateau basalt descends to ∼540 m a.s.l.(Fig. 8), and along its NE margin it is quite altered,suggesting contact with the Gediz at its contem-poraneous level (Westaway et al., 2004). As Fig. 8 illus-trates, this altitude limit is noticeably lower than the∼560 m limit of the Sarnıç plateau basalt aroundÇakırca, ∼3.5 km farther east (site 4 in Fig. 7a), more

the principal hilltop summits in the Ulubey Formation limestone, withm SRTM data: 1, unnamed, N740/N735 m; 2, Sakaryol Tepe, 745 /düzü Sırtı, N765 m, N750 m; 6, Kocakır Tepe, 733 /N725 m.

Fig. 11. Similar imagery to Fig. 8 but for the vicinity of Dereköy. Site Y denotes İnkale Tepe (height 584 m on HGK map, N580 m by SRTM); siteβ denotes Çilo Tepe (height N565 m on HGK map, N535 m by SRTM).



so than can be explained by any plausible westwardGediz palaeo-gradient, suggesting that the İbrahimağaplateau basalt post-dates the Sarnıç plateau basalt. Ourdating of sample 01TR97 to 999±21 ka confirms this.

3. Uplift modelling

3.1. Altitude measurements

In order tomodel incision as a proxy for surface uplift,accurate measurements of altitudes of fluvial terracesand of present river levels is an obvious requirement. Theprincipal source of topographic information for thisstudy region is the series of 1 :25,000 scale topographicmaps prepared by the Turkish Army General Commandof Mapping (Harita Genel Komutanlığı or HGK). Thesemaps have 10 m contours, with supplementary 5 mcontours in areas of low relief, and show spot heights tothe nearest 1 or 0.1 m, shading being used to indicate thepositions of basalt cliffs.Westaway et al. (2004) used thissource of information, supplemented by ground truthingwith a handheld GPS receiver, to estimate the level of theGediz and the heights of many of its terrace deposits to anominal precision of ∼5–10 m. Maddy et al. (2005, inpress) used a Total Station to survey the heights of manyGediz deposits, relative to a datum at the level of the riverwhere it is crossed by the Ankara-İzmir highway nearKale Tepe (Fig. 7a). These altitude measurements, to anominal precision of 0.1 m, were used by Maddy et al.(2005, in press) to construct longitudinal and transverseterrace profiles (Fig. 7a,b), although the raw data havenot been disclosed. The datum used by Maddy et al.(2005, in press) is not a HGK benchmark for which aspot height is available, so the absolute heights of each oftheir survey points have not previously been established.However, according to the local HGK map (1 :25,000scale sheet Uşak K21-c4), this datum is very close to400 m a.s.l.. We thus add 400 m to convert the relativeheights reported by Maddy et al. (2005, in press) toabsolute heights, for comparison with other evidence.

A third source of topographic information is nowavailable, from the Shuttle Radar Topographic Mission(SRTM), as has been used to generate Figs. 8–11. This isbased on data collected using an interferometric radaraltimeter, which operated from the space shuttle Endea-vour during 11–21 February 2000 (see, e.g., Rabus et al.,2003, or Westaway et al., 2006, and references therein,for more details). The SRTM provides measurements ofa representative value of the surface altitude within a gridof 90×90 m ‘footprints’. At conical summits, such asÇilo Tepe (β in Fig. 11), one expects this technique tounderestimate summit altitude, because the ‘footprint’

closest to the summit will also sample the hillsides, as isindeed so (see Fig. 11 caption). However, whenmeasuring heights of basalt plateaus, where there istypically negligible change in altitude across eachinterval of 90 m, one should expect this technique toproduce accurate heights. Nonetheless, for the set ofrepresentative heights investigated in Figs. 8–11, theSRTM altitudes typically underestimate the HGKaltitudes by ∼10±5 m. We do not understand the sourceof this discrepancy: a systematic mismatch of 10 mwould imply use of a different zero level for altitude, butthere also seems to be a measurable random componentin these relative heights, indicating some other source oferror in one or both of the sets of determinations. Similarproblems—of similar magnitude—with SRTM datahave been noted elsewhere in the eastern Mediterraneanregion by Miliaresis and Paraschou (2005).

To illustrate the potential importance of uncertainty inheight, we consider as an example the altitude of theGediz gravel beneath the basalt that yielded dated sample00YM23 at Çakırca, at the western margin of the SarnıçPlateau (Fig. 8). Westaway et al. (2004) located this siteon a HGK map and thus deduced that this gravel is∼560m a.s.l. Using the GPS fix reported byWestaway etal. (2004) we have located this site on the SRTM imagein Fig. 8, deducing instead an altitude of ∼555 m a.s.l.However, this GPS fix may itself be in error, because itwas of necessity taken near the base of a basalt cliff,where visibility of the sky was limited. Finally, the TotalStation survey by Maddy et al. (2005) placed this point(their site 4; Fig. 7a) ∼150 m above datum, or at anestimated∼550m a.s.l. By analogy, similar uncertaintiesof the order of∼5–10m can be expected in both absoluteand relative altitudes of other sites; the uplift modellingpresented below, which uses a combination of heightsderived from HGK maps and from Maddy et al. (2005),is thus subject to such uncertainty.

3.2. Age constraints

The new dating evidence helps considerably toconstrain the uplift chronology of the study region, aswell as modifying some of the existing age constraintsused in the previous modelling by Westaway et al.(2004). First, the very young age now favoured for theβ4 basalt flows means that the depths to which they havebeen incised have no bearing on the long time-scaleregional uplift history; they simply represent recovery bythe Gediz of a roughly equilibrium gradient profile afterits course was temporarily dammed by basalt. This evi-dence that the river has been able to recover from dam-ming so quickly supports the deduction by Westaway et


al. (2004) that for most of the time since the Kulavolcanism began it has maintained a roughly equilibriumdownstream profile, such that the heights of its terracesreflect regional uplift, not local disturbances caused bybasalt damming. The coarseness of the modern bedloadof the Gediz (evident in Figs. 3d and 5b–c) is presumablyconducive to this rapid re-incision. Second, our dates forthe β3 basalts indicate ∼40 m of incision since 175 ka atPalankaya (time-averaged rate 0.23 mm a−1), ∼25 m ofincision since 180 ka at Kula Bridge (0.14 mm a−1),and ∼35 m of incision since 236 ka upstream of KulaBridge at Dereköy (0.15 mm a−1). This evidence sug-gests that the rate of incision—and thus, uplift—in-creases substantially downstream between Kula Bridgeand Palankaya.

Regarding the β2 basalt, Westaway et al. (2004)estimated 205 m of incision by the Gediz since eruptionof the İbrahimağa plateau basalt, lowering its level from∼540 m to its present ∼235 m a.s.l. beside the westernpart of this plateau. The dating of our sample 01TR97indicates that this incision has occurred at a time-averaged rate of 0.21 mm a−1.

For the Sarnıç Plateau, we now have three quitetightly grouped dates (all ±2σ), of 1264±30 ka(00YM23), 1219±27 ka (01TR74) and 1174±26 ka(00YM22). If it is inferred that these three samples are allfrom the same flow unit, like at Palankaya, then theyindicate an overall weighted mean age for it of 1215±16 ka (±2σ). Following Westaway et al. (2004), weestimate the altitude of the Gediz gravel underlying thisbasalt at Çakırca as ∼560 m a.s.l. (see above), ∼185 mabove the present river level, indicating a time-averagedincision rate since 1215 ka of ∼0.15 mm a−1. Thiscalculation assumes the differences between these threedates are due to random error.

However, another possibility is that the SUERC datefor sample 00YM23 is older than the LSCE dates for thesame reason as the SUERC dates for the β4 categorybasalts are older (see above), this reason possibly relat-ing to the procedure used at SUERC to correct foratmospheric argon isotope ratios in the laboratory. If so,then this difference is a form of systematic error, makingit inappropriate to average the three dates. Averaging thetwo LSCE dates instead produces a weighted mean ageof 1196±19 ka (±2σ). The subsequent time-averagedrate for 185 m of incision is still ∼0.15 mm a−1, andwould remain so if the lower height of ∼550 m a.s.l.,derived from Maddy et al. (2005), were used instead.

For the Burgaz Plateau, we now have two dates, of1186±156 ka (Ar–Ar; ±2σ; based on Richardson-Bunbury, 1996) and 1014±23 ka (±2σ; 00YM25). Evenallowing for the large error margin in this Ar–Ar date,

these two numerical ages are barely concordant at the±2σ level (i.e., at a 95% confidence limit), suggestingthe presence of a systematic error. Given the individualages of the two sample splits that contributed to this Ar–Ar date, of 1370±200 and 1120±120 ka (both ±2σ), wethus consider it probable that the older one was con-taminated by inherited argon. The weighted mean age ofthe younger of these Ar–Ar splits and our date is 1018±23 ka, not significantly different from the age deter-mined for our own sample. As already noted, in thevicinity of the SE margin of the Burgaz Plateau at KaleTepe, the Gediz has incised since this basalt eruptionfrom ∼540 to ∼400 m a.s.l. or by ∼140 m, at a time-averaged rate of 0.14 mm a−1.

Westaway et al. (2004) inferred that both the Burgazand Sarnıç plateau basalts erupted at ∼1200 ka, andattributed the lesser amount of subsequent incision ofthe former than of the latter to a lateral variation in upliftrates. It now seems clear instead that the lesser incisionof the Burgaz plateau basalt is mainly a consequence ofit being younger than the Sarnıç plateau basalt. In con-trast, there is evidence that rates of incision, and thus,uplift, increase west of the Sarnıç Plateau, as indicatedby the faster incision of the basalt flanking the north-western part of the İbrahimağa Plateau and at Palankaya.

Regarding the total incision since the Pliocene,Westaway et al. (2004) noted that in the vicinity of theİbrahimağa Plateau (Fig. 10) the top of the UlubeyFormation has been incised by ∼410 m, from 745 to∼335 m a.s.l. Farther upstream within the Selendi Basin,the Ulubey Formation limestone is nowhere preservedadjacent to the Gediz or the β2 basalts. Outliers of it,typically tens of metres thick, cap the highest land in thecentral-southern Selendi Basin north of the Gediz,reaching 890 m a.s.l. at b, 805 m at c, 865 m at d, and825 m at e in Fig. 2 (Westaway et al., 2004). After aninterval of several kilometres within the badland land-scape in the Ahmetler Formation tuffites and sands, thislimestone is again observed farther east, where it is∼100 m thick, its upper surface being ∼920 m a.s.l.north of Ulucak (Westaway et al., 2004; Fig. 1) and at∼880 m west of this village, ∼5 km farther SW.Westaway et al. (2004) extrapolated the gentle SW tilt ofthis limestone surface from the Ulucak area to ∼840 ma.s.l. in the vicinity of Ziftçi Tepe (O in Fig. 2), east ofthe Burgaz Plateau. This hill (summit, 618 m), is cappedby a ∼750 m (E–W) by ∼400 m (N–S) expanse of β2basalt (which remains undated; it erupted from neck 76,Delihasan Tepe, to the SW; V in Fig. 2). This basaltsurface is above 610 m a.s.l., overlying the top of theAhmetler Formation at ∼560-570 m a.s.l., ∼150 mabove the Gediz at∼415m a.s.l. (Westaway et al., 2004).


A total of∼425 m of incision is thus indicated here, withalmost two thirds (∼275 m) pre-dating the DelihasanTepe eruption and the remaining third (∼150 m) post-dating it. We thus estimate that the incision below thetop of the Ulubey Formation limestone, projected intothe vicinity of the Burgaz and Sarnıç plateaus, has beenbetween these limiting values of ∼410 and ∼425 m.

Regarding MIS assignments, it is thought that thatwithin the eastern Mediterranean region—like in west-ern and central Europe—river terraces typically formduring even-numbered cold stages of the oxygen isotopetimescale, being associated with increased sedimentfluxes caused by reductions in vegetation cover (e.g.,Collier et al., 2000; Macklin et al., 2002). Maddy et al.(2005) inferred that Gediz high terrace I formed duringthe MIS 38 cold stage at ∼1240 ka (or during the MIS38-37 warming limb). This was based on taking the ageof the Burgaz Plateau basalt as 1245 ka, determined asthe unweighted mean of the sample splits for Richard-son-Bunbury's (1996) Ar–Ar date. As already noted, useof the weighted mean of these splits would give an age of1186 ka, from which one could infer that terrace I datesfrom MIS 36 at ∼1200 ka. However, our revised datingmakes this Burgaz Plateau basalt considerably younger,and our preferred age for it of 1014±23 ka suggestsdeposition of the gravel of Gediz high terrace I duringMIS 28 (∼1020 ka). If it is further assumed, followingMaddy et al. (2005), that the Gediz high terraces beneaththe Burgaz Plateau formed in response to a sequence ofconsecutive ∼40 ka cold stages, then the oldest (XI) canbe assigned to MIS 48 at ∼1440 ka (Fig. 7b).

In contrast, the dating evidence now available sug-gests that Gediz high terrace I at Çakırca pre-dates theeruption of the Sarnıç plateau basalt at 1215±16 ka, sothis gravel may well have aggraded in MIS 38(∼1240 ka), fortuitously as suggested by Maddy et al.(2005). However, such an interpretation would implythat Gediz high terrace I at Çakırca formed at the sametime as Gediz high terrace VI in the Burgaz Plateau. Assurveyed by Maddy et al. (2005), this terrace in theBurgaz Plateau has since been incised by 173 m, con-sistent with the ∼175–185 m of incision estimated atÇakırca (see above). The five consecutive higherterraces (III–VII) interpreted by Maddy et al. (2005)in the Sarnıç Plateau may thus represent the same coldstages as terraces VII–XI in the Burgaz Plateau. Maddyet al. (2005) reported Gediz high terrace X at two sites inthe Sarnıç Plateau, their sites 45 and 46 (Fig. 7a) in thevertical cliff face adjoining Toytepe neck (the localityillustrated in Figs 14a and b of Westaway et al., 2004),recognising thin Gediz gravel above the BalçıklıdereMember fluvial sand and below the predominant lime-

stone tributary (fan) gravel in this area. However, asnoted above, another possibility is that the best date forthe Gediz gravel at Çakırca is 1196±19 ka, in whichcase this gravel could have formed either in MIS 38 or36. Each of the older deposits of Gediz gravel beneaththe basalt of the Sarnıç Plateau may likewise possibly beone 40 ka climate cycle younger than estimated above.

3.3. Physical properties of the crust

Previous experience (e.g., Westaway, 2001, 2002a,b,2004a; Westaway et al., 2002) indicates that isostaticuplift responses revealed by river terrace staircases canbe modelled as a function of rates and spatial scales ofsurface processes occurring in the region and of physicalproperties of the underlying continental crust. Assumingthat erosion and sedimentation are driving the verticalcrustal motions (cf. Westaway, 2002c; Westaway et al.,2004), relevant parameters quantifying rates of surfaceprocesses include the local spatial-average erosion rateU, the characteristic horizontal scale Le over which thiserosion occurs, and the corresponding scale Ls and spa-tial-average rate of sedimentation in the adjacentdepocentre to which this eroded material is transported.Relevant crustal parameters include the densities ofcontinental crust, mantle lithosphere, asthenosphere,and any water that loads the depocentre (as an approx-imation, the sediment load is considered part of thecrustal rock column), respectively, ρc, ρm, ρa, and ρw,and the thermal diffusivity of the crust, κ. The isostaticresponse is influenced by the thickness of the mobilelower-crustal layer, W, which equals the differencebetween the crustal thickness zm and the thickness of theupper-crustal brittle layer zb; W determines the rate atwhich material can flow within the lower-crust in re-sponse to a given pressure gradient P resulting from agiven lateral variation in zb. The depth zb corresponds toa particular temperature in the region of ∼350 °C, and isthus determined by the geothermal gradient in the crust,taking account of perturbations to this gradient by thesurface processes that are occurring. The uplift responsewill also be affected by the extent to which loading orunloading of the mantle lithosphere causes flexure inthis layer, which will depend on the thickness of themantle lithosphere and its mechanical properties (whichdetermine its flexural wavelength, λ) as well as on therelative magnitudes of λ and the characteristic scales ofthe surface processes, Le and Ls. Crustal extension willof course also affect the state of isostatic equilibrium. Itseffects are neglected in the present modelling; althoughthe crust of western Turkey is extending rapidly overall,as already noted there is no significant extension of the


brittle upper crust in the vicinity of the reach of theGediz around Kula (cf. Westaway et al., 2004).

Regarding observational constraints, Westaway(1994) deduced from the sedimentation rate in the east-

Fig. 12. Uplift results using the technique of Westaway (2001). Observationapredicted uplift history are listed in Table 2. See text for discussion.

ern Aegean Sea a spatially averaged and time-averagederosion rate throughout the Gediz drainage catchment of∼0.1 mm a−1 during the Middle to Late Pleistocene.Some material eroded within this catchment has not

l data points are labelled. The model parameters used to generate each

Table 2Model parameters for initial uplift modelling

Parameter Units Value

Prediction 1 Prediction 2 Prediction 3

Assumed valuesc °C km−1 20 20 20κ mm2 s−1 1.2 1.2 1.2Wi km 17 16 15to1 ka 18000 18000 18000ΔTe1 °C −20 −20 −20to2 ka 3100 3100 3100ΔTe2 °C −27.5 −23.5 −21to3 ka 900 900 900ΔTe3 °C 0 −8 −8

Derived valueszb km 11.1 12.2 13.3Tm °C 727.8 705.6 683.3U (875 ka) m 118 125 122U (950 ka) m 130 135 132U (3100 ka) m 411 410 411vumax mm a−1 0.182 0.175 0.177t(vumax) ka 1800 1950 2100

Parameters have the following meaning: c, geothermal gradient in thelower crust; κ, thermal diffusivity of the crust; Wi, paramaterquantifying the difference in depth between the base of the brittlelayer (zb) and the level at whoich lower-crustal flow is mostconcentrated; to and ΔTe, timing and magnitude of each phase ofLCFF, with subscripts 1, 2 and 3 indicating individual phases; zb, theestimated depth of the base of the brittle upper crust (calculated as zm−Wi / 0.9 (see main text; zm being the Moho depth, taken as 30km); Tm,the estimated Moho temperature (calculated as 350 °C+(zm−zb)×c);U(t), the predicted amount of uplift since time t; vumax the maximumpredicted uplift rate and t(vumax) its timing.


reached the sea, being instead deposited in the AlaşehirGraben (Fig. 1), so this estimate is a lower bound to theerosion rate in the part of this catchment that is eroding.An upper bound to the time-averaged erosion rate in thepresent study region is ∼0.13–0.14 mm a−1, corre-sponding to ∼410–425 m of erosion (see earlier discus-sion) since ∼3 Ma, but some parts of this region (e.g.,the areas of in situ Ulubey Formation limestone and theβ2 basalt plateaus; Fig. 2) have clearly eroded sub-stantially less than this. Fission-track dating of the meta-morphic basement that has been exhumed by erosion inthis area (e.g., Ring et al., 2003) provides no relevant databecause the amounts of erosion since the Pliocene havebeen so small.

The crustal thickness in the vicinity of Kula hasbeen determined as 30±1 km from detailed studies ofseismic wave propagation (Saunders et al., 1998). Wethus adopt a nominal value of zm=30 km. The regionalaverage heat flow in western Turkey is high, being∼110 mW m−2 (Ilkışık, 1995). It is particularly higharound Kula, where analysis of the Acısu hot spring(M in Figs. 2 and 8) indicates a local value of 123±32 mW m−2 (Ilkışık, 1995), requiring a geothermalgradient in the uppermost crust of ∼40 °C km−1. Ilkışık(1995) estimated that about half this heat flow arisesthrough radiogenic heat production in the upper crust(i.e., Qr∼55 mW m−2), the rest being the result ofconduction from the Earth's interior. Standard theory(e.g., Lachenbruch, 1970), which ignores any perturba-tion to the geotherm caused by erosion (cf. Westaway,2002c), indicates that Qr =YoD where Yo is the near-surface radiogenic heat production and D is the depthscale over which radiogenic heat production decreasesto 1 /e of its near-surface value. However, neither Yo norD are well-constrained for western Turkey. Nonetheless,using this theory (e.g., Lachenbruch, 1970) one canpredict that in this region the depth of the base ofthe brittle layer, zb, corresponding to a temperature of∼350 °C, is ∼10–14 km (taking D=2.5 km andYo=22 μW m−3 would give zb∼10.3 km, whereas tak-ing D=14 km and Yo =3 μW m− 3 would givezb∼13.5 km). The depth-limit of seismicity indepen-dently suggests a relatively shallow base of the brittlelayer, ∼10 km deep (e.g., Jackson and McKenzie,1988). Saunders et al. (1998) also inferred that theseismic S-wave velocity in the Kula area decreasesbelow a depth of 13 km, providing a third estimate of zb.

Using the same range of values as above (for Yo andD) places the base of the lithosphere (at a temperature of∼1400 °C) at a depth of ∼53–59 km, implying that themantle lithosphere in this region is b∼30 km thick (cf.Ilkışık, 1995; Westaway, 2006). Taking account of the

expected perturbation to the geotherm caused by theregional erosion that typefies western Turkey, the base ofthe lithosphere is expected somewhat deeper that thesesimple calculations predict (cf. Westaway, 2002c), fromwhich it can be inferred that theMoho temperature in thisregion (at a depth of 30 km) is b∼700 °C. However,rather than assuming any particular Moho temperature,the modelling will consider a range of possible values.The remaining model parameters will either be assignedfixed nominal values or be allowed to vary in order tomatch the evidence.

3.4. Modelling techniques and results

These uplift data will be modelled using two tech-niques. The first, by Westaway (2001; Westaway et al.,2002), calculates the isostatic response to phases oflower-crustal flow forcing (LCFF) induced by cyclicloading of the Earth’s surface, such as may be caused bygrowth and decay of local ice sheets and fluctuations inthe water load applied to the continental shelf as a result


of eustatic sea-level variations. For this technique, thecharacteristic timescale of the uplift response inducedfollowing the start of a phase of LCFF depends on theparameter Wi= zi− zb, zi being the depth at which theinduced lower-crustal flow is concentrated (Westaway,2001; Westaway et al., 2002). For the viscosity para-meterisation used for the lower crust, with a linearviscous rheology dependent only on temperature, and foran assumed uniform steady-state geothermal gradient inthe lower crust, the depth zi is roughly 9/10 of the waydown from zb to the depth zm of the Moho (Westaway,1998). If zm is known in a region, one may thus estimatezb as zm—10 Wi / 9, for comparison with observations.

As in the Westaway et al. (2004) study, this techniqueis applied to western Turkey as an approximation, itbeing clear that the uplift of this region is caused byincreased rates of erosion and not by cyclic surfaceloading. However, tests by Westaway (2002b) showedthat predicted uplift histories assuming cyclic surfaceloading and assuming a non-steady-state response tochanging rates of erosion can be very similar to eachother, justifying the use of the Westaway (2001) tech-nique as an approximation. The Westaway (2001) tech-nique is also easier to apply, requiring fewer modelparameters to be constrained, and also lends itself tosolutions involving multiple phases of LCFF.

The second technique, by Westaway (2002c), calcu-lates the non-steady-state isostatic uplift response to aregion that experienced an increase in erosion rates (or aswitch from sedimentation to erosion) at time to, as-suming that the region adjoins a depocentre in which theeroded material is re-deposited. In such a situation, thegeothermal gradient is perturbed in a predictable mannerin both the eroding sediment source region and thedepocentre, causing the base of the brittle layer to advect(relative to the level of the eroding land surface) upwardbeneath the sediment source and downward beneath thedepocentre. The resulting lateral variation in its depthwillcreate a lateral pressure gradient, which will drive mobilelower crust from beneath the depocentre to beneath thesediment source. The technique is more difficult to applyas it requires more model parameters to be constrainedthan the alternative, and it can also only be used to modela single phase of LCFF caused by an increase in erosionrates; the more general problem involving multiplephases has no closed analytic solution.

Fig. 13. First set of uplift results using the technique of Westaway (2002c). Obwhether the total uplift since ∼3 Ma is ∼410 or ∼425 m. To avoid clutter, oGediz in the Kula area are shown. The model parameters used to generate eachbeen determined using a nominal water density of 0, indicating that subaerialdiscussion.

Since its previous use by Westaway (2002c) andWestaway et al. (2004), this computer program has beenmodified. In these previous studies, it solved equation(B19) of Westaway (2002c) to obtain at each time stepthe depth zb corresponding to the nominal 300 °C tem-perature at the base of the brittle upper crust. In suchcalculations, a constant nominal crustal thickness za wasassumed. Calculations of inflow and outflow of lowercrust, and of the associated isostatic response, werehandled separately. This simplification was consideredappropriate, because the model crustal thickness seldomchanges by more than∼10 m in any 25 ka time step. Theprogram now calculates zb at each time step for thenominal crustal thickness, as before, then scales it by theratio of current crustal thickness to nominal crustalthickness. This also involves making an approximation,but is considered more realistic than before. The dif-ference between methods is particularly important whenmodelling cumulative effects of lower-crustal flow overrelatively long timescales, as in the present study, orwhen modelling relatively rapid uplift, accompanied byvigorous lower-crustal flow. Compared with the oldversion, the new algorithm results in smaller-magnitudevariations in zb, thus requiring a lower effective viscosityηe for the lower crust than before in order to sustain agiven geometry of uplift.

This use of a 300 °C temperature threshold for thebase of the brittle upper crust, rather than the morerealistic ∼350 °C value mentioned earlier, enables zb tobe determined at appropriate values of ∼13 km, eventhough the computer program used does not take accountof the upward increase in the geothermal gradient causedby radiogenic heating at shallow depths. A key require-ment in this type of modelling is to base calculations onappropriate values for the thickness W of the lower-crustal layer, since fluxes of lower-crust are highlysensitive to this parameter, being dependent onW3 (e.g.,Westaway, 2002c). Use of this particular approximationin lieu of more complicated calculations that incorporateradiogenic heating in the upper crust was justified as aresult of numerical tests by Westaway (2002c).

It has been observed that, to a first approximation, theGediz terraces capped by the β2 basalts in the Burgazand Sarnıç Plateaus and the β3 basalts in the reach of themodern gorge between Dereköy and Kula Bridge allsupport uplift rates of ∼0.14–0.15 mm a−1. This view

servational data points are shown, subject to uncertainty depending onnly the youngest and oldest terraces of the high terrace staircase of thepredicted uplift history are listed in Table 3. Each of these solutions hasdeposition has been assumed throughout each model run. See text for


differs from the interpretation proposed by Westaway etal. (2004), who deduced a significant increase in upliftrates in the Middle Pleistocene. However, it can now beseen that that earlier interpretation was incorrect; it wasbased on dating of the β3 and β4 basalts in the vicinityof Kula Bridge that is now known to be invalid, and onthe assumption of the same incision rate at Palankaya asat Kula Bridge.

Using the Westaway (2001) technique to model arecord of uplift like this, with a relatively uniform rate,requires the associated LCFF to have begun well beforethe observational record and to be associated with arelatively thick lower-crustal layer. Prediction 1 inFig. 12 shows one such solution, constrained to match∼410 m of uplift since 3.1 Ma. It fits the Gediz highterrace staircase very well, consistent with it dating fromconsecutive 40 ka cold stages from MIS 48 to MIS 28. Italso fits the data from Dereköy and Kula Bridge rea-sonably well, although slightly underestimating the in-cision documented at each of these localities. Thissolution requires a rather thick lower-crustal layer,consistent with zb∼11 km, too shallow to match the dataof Saunders et al. (1998), and a Moho temperature of∼730 °C, which is probably too high (Table 2).

In many localities worldwide, uplift rates are ob-served to have increased in the Middle Pleistocene (e.g.,Kukla, 1975, 1978; Westaway, 2001, 2002a; Bridglandand Westaway, in press), an effect that can be attributedto a consequence of more vigorous LCFF beginning at∼0.9 Ma, perhaps accompanying systematic increases inerosion rates during the major cold stages that beganaround this time when the Milankovitch forcing switch-ed from characteristic ∼40 to ∼100 ka periodicity (e.g.,Mudelsee and Schulz, 1997). The combination of therelatively high uplift rates already occurring in the Kulaarea in the Early Pleistocene and a relatively thick lower-crustal layer (consistent with the high local heat flow)will have resulted in this Middle Pleistocene effect beingrelatively subdued and only becoming significant after arelatively long time lag following the inferred start of itsLCFF at∼0.9 Ma (cf. Westaway, 2002b). Prediction 2 inFig. 12 shows an alternative uplift solution, whichincludes an additional phase of LCFF with this timing.This fits the late Middle Pleistocene data better thansolution 1, while maintaining an equivalent fit to thatgiven before to the Gediz high terrace staircase. Theslightly greater inferred zb of ∼12 km and the resulting

Fig. 14. Second set of uplift results, using the technique and data as Fig. 13. Tlisted in Table 3. Each of these solutions has been determined using the samwater density of 1000 kg m−3 has been assigned, indicating that marine ddiscussion.

lower estimated Moho temperature of just over 700 °Cseem more reasonable than for solution 1.

The regionmost similar to this part of western Turkey,in terms of crustal properties, that has been modelled todate, is the terrace staircase of the River Allier in thevicinity of the Perrier Plateau south of Clermont-Ferrandin central France (e.g., Pastre, 2004), where the crust is∼30 km thick and has a surface heat flow of ∼110 mWm−2 (Westaway, 2004a). Modelling of the well-datedterrace staircase in this area required Wi 15 km, so itseems reasonable to consider the consequences ofassuming such a value for Kula, as represented by pre-diction 3 in Fig. 12. This solution can once again give anexcellent fit to the Middle Pleistocene data and to the∼410 m of total incision, but it is no longer possible to fitthe Gediz high terrace staircase to a series of consecutive40 ka cold stages. It is instead estimated that the oldestterrace (XI) dates from MIS 52, indicating that two coldstages have been ‘missed’ between that and the youngestfrom MIS 28. Possible cold stages that may have beenmissed (but see also below) could be MIS 50 (betweenterraces XI and X) andMIS 40 (between terraces VII andVI), which would correspond to two of the highestobserved terrace bluffs (Fig. 7b). This revised solutionpredicts zb ∼13 km, consistent with the data fromSaunders et al. (1998), and gives a realistic Mohotemperature of ∼680 °C (Table 2).

The alternative modelling using the technique ofWestaway (2002c) assumes a sudden increase in erosionrates at 3100 ka to a value that remains constant there-after. As already noted, this technique cannot investigatethe consequences of any additional subsequent increasesin erosion rates, such as may have occurred around∼900 ka. By analogywith solutions 2 and 3 in Fig. 12, itspredictions are thus expected to slightly underestimatethe total uplift since 3100 ka and the uplift ratesdetermined from the amounts of incision at Dereköyand Kula Bridge.

Fig. 13 illustrates a family of solutions in which a200 km wide eroding sediment source region is coupledto a depocentre, also 200 km wide, via a 100 km widehinge zone. Fig. 22 of Westaway et al. (2004) shows adiagramatic representation of such a model. The modelsediment source region is intended to represent the reachof the Gediz around and upstream of Kula (Fig. 1),whereas the model depocentre is intended to representwhat is now the eastern part of Aegean Sea, underlain by

he model parameters used to generate each predicted uplift history aree model parameters as the corresponding solution in Fig. 13, except aeposition has been assumed throughout each model run. See text for


a stacked sedimentary sequence comprising material thathas been transported by the Gediz (and other rivers in theregion). The Aegean Sea is well known to have becomesignificantly wider and deeper during the Quaternary(e.g., Papanikolaou, 1993), consistent with net crustalthinning accompanying outflow of lower crust, as isrequired to sustain the observed uplift of the adjacentonshore regions (e.g., Westaway et al., 2004). To incor-porate the fact that much of this sea area did not exist formuch of the time span covered by the model, for thesolutions in Fig. 13 the density of the model seawatercovering the depocentre has been set to zero, implyingsubaerial deposition. This clearly misrepresents the mostrecent part of the record, during which the water depthhas increased to its present∼1000 mmaximum, and thusunderestimates the associated amount of loading of thecrust, introducing a number of systematic errors into themodel. Conversely, the solutions in Fig. 14 assumeseawater loading throughout, which will misrepresentthe early part of themodel andmake the predicted overalleffect of water-loading greater than will have occurred inreality. These two sets of solutions can thus be regardedas end-members of a range of possibilities, with the ‘true’solution lying somewhere in between.

These solutions also assume overall Airy isostaticcompensation; given the high Moho temperatures ex-pected, the mantle lithosphere is assumed to be suf-ficiently thin and weak so that it will flex on the scales ofthe loads that are applied in the model. Regarding theuplifting sediment source region, the associated devel-opment of a lower-crustal ‘root’ will mean that the pre-dicted rate of increase in altitude of the eroding landsurface will be Fa= (ρa−ρc) /ρa times the predicted rateof crustal thickening. The predicted uplift rate of markersthat are not themselves eroding (such as Gediz terracescapped by basalt in the Burgaz Plateau) will exceed thispredicted rate of altitude increase by an amount equal tothe assumed spatial average erosion rate of the landsurface as a whole (e.g., Westaway, 2002c). In thismodelling, the effective viscosity ηe of the lower crust(the viscosity that an isoviscous layer would need to havein order to produce the same overall flow rate—or thesame average flow velocity vx—in response to a givenapplied pressure gradient as is observed in a real lower-crustal layer with temperature dependent viscosity;Westaway, 1998), is specified in advance for eachsolution. The associated Moho temperature is estimatedusing the ‘rule of thumb’, fromWestaway (1998) that theviscosity at the Moho is ∼1/60 ηe. The Mohotemperature is then estimated for this Moho viscosityassuming the preferred viscosity–temperature calibra-tion from Westaway (1998).

The solutions in Fig. 13 assume values ofηe of 5 (M1),3 (M2), 2 (M3) and 1 (M4) ×1018 Pa s, consistent withMoho temperatures in the range∼630–680 °C (Table 3).As indicated in Fig. 13, none of these solutions canpredict even the 270m lower bound to the uplift observedbetween 3100 ka and the time of formation of Gediz highterrace I. The main reason why such solutions will un-derestimate uplift is failure to incorporate water-loadingof the depocentre. It thus follows that such loading wasalready appreciable by the late Early Pleistocene, if notearlier, so we turn our attention to the solutions in Fig. 14that assume such water loading throughout.

Solution M3 in Fig. 14 fits well the lower bound touplift indicated by Gediz high terrace I. It also fits wellthe lower bound for Gediz high terrace XI, with thisterrace assigned to MIS 52 (see above), less so if it isfitted to MIS 48. Likewise, solution M4 fits the upperbounds to uplift estimated for these terraces. However,solution M4 predicts almost ∼1200 m of bathymetry inthe adjacent part of the Aegean Sea, which exceeds itstypical observed depth, whereas solution M3 predictesb1000 m of water depth, which is more appropriate. Wethus regard solution M3 in Fig. 14 and Table 3 as ourpreferred solution.

All solutions in Fig. 14 underestimate the total upliftfrom 3.1 Ma to the present-day (Figs. 13a, 14a) and alsopredict lower rates of uplift and incision (Figs. 13b, 14b)than are observed for those Gediz terraces capped by β3basalt. We presume this to be a consequence of an in-crease in uplift rates in the Middle Pleistocene, as inFig. 12, caused by an increase in erosion rates in the lateEarly Pleistocene to a value slightly above 0.1 mm a−1.We note that for our preferred solution (solution M3 inFig. 14) the predicted uplift response at 2200 ka, 900after the assumed acceleration in uplift, is 0.1249 mma−1, ∼30% more than the assumed erosion rate of0.1 mm a−1 that is forcing this uplift. We thus deduce byanalogy that the incision rates of ∼0.14–0.15 mm a−1

observed since the late Middle Pleistocene in the KulaBridge/Dereköy areas are ∼30% faster than the typicalpost-0.9 Ma erosion rate, enabling this rate to be esti-mated as∼0.11–0.12mm a−1. However, as noted above,the technique used cannot model the isostatic conse-quences of a succession of changes in erosion rates, sothis estimate cannot at present be tested.

4. Discussion

The significant improvement in age control providedhere (Table 1) means that, finally, the history of volca-nism and fluvial deposition and incision in the Kularegion can be considered well-constrained. This


chronology differs substantially from those proposed inother recent studies, notably by Westaway et al. (2004)and Maddy et al. (2005), which did not have the benefitof this improved set of dates. The principal mistakes nowevident in the Westaway et al. (2004) study were itsfailure to recognise lateral variations in incision since thelateMiddle to Late Pleistocene along the studied reach of

Table 3Parameters used in modelling of uplift in response to erosion

Parameter Units Value

M1 M2 M3 M4

Assumed valuesHco km 30 30 30 30Hmo km 70 70 70 70Tmo °C 700 700 700 700to ka 3100 3100 3100 3100Uoe mm a-1 10-9 10-9 10-9 10-9

Ue mm a-1 0.1 0.1 0.1 0.1Uos mm a-1 10-9 10-9 10-9 10-9

Us mm a-1 -0.1 -0.1 -0.1 -0.1Le km 200 200 200 200Lh km 100 100 100 100Ls km 200 200 200 200Fa 0.129 0.129 0.129 0.129zwo m 0 0 0 0Ed Ma-1 0 0 0 0Em Ma-1 0 0 0 0ηe 1018

Pa s5 3 2 1

Derived valuesTm °C 626.1 643.6 657.9 683.6ηm 1016

Pa s8.333 5.000 3.333 1.667

Predictions for ρw=1000 kg m-3

zw m 615.2 826.8 983.0 1168.8vu mm a-1 0.1227 0.1264 0.1271 0.1248Y m 364.6 383.2 396.9 413.1Hce km 30.5225 30.7010 30.8323 30.6277Hcs km 29.4775 29.2990 29.1677 29.3443zbe km 12.8448 12.7862 12.7289 12.6612zbe km 12.8641 12.9199 12.9751 13.0404Δzb km 19.3 133.7 246.2 379.2ηe /Δzb 1015

Pa sm-1259.1 22.44 8.123 2.637

Predictions for ρw=0 kg m−3

zw m 375.0 445.2 482.2 515.3vu mm a-1 0.1163 0.1158 0.1147 0.1057Y m 358.8 367.8 372.6 376.8Hce km 30.3781 30.6092 30.4851 30.5179Hcs km 29.6219 29.5517 29.5149 29.4821zbe km 12.7840 12.8135 12.8290 12.8428zbe km 12.9271 12.8965 12.8804 12.8661Δzb km 143.4 83.0 51.4 23.3ηe /Δzb 1015 Pa

s m-134.87 36.14 38.91 42.92

Notes to Table 3:Hco and Hmo are the initial thicknesses of the crust and mantle

the Gediz and the misinterpretations resulting from thesystematic errors in the K–Ar dating of the young basaltsamples. Both Westaway et al. (2004) and Maddy et al.(2005) drew incorrect conclusions about the Gediz ter-races capped by β2 basalts, not least by assuming thatthese basalts all have the same age. The evidence nowavailable that the Burgaz plateau basalt is about twohundred thousand years younger than the Sarnıç plateaubasalt also means that theMaddy et al. (2005) correlationscheme for the Gediz high terrace staircase (Fig. 7a)breaks down between these plateaus. However, the keyconclusion of Maddy et al. (2005), that the staircases ofGediz high terraces preserved beneath these plateausrepresent a succession of ∼40 ka Milankovitch cycles inthe Early Pleistocene, is strongly supported by the pre-sent uplift modelling.

Of the six β2-category volcanic necks in the Kulavolcanic field (Fig. 2), at least two (Tavşan Tepe, neck58, and Delihasan Tepe, neck 76), and possibly a third(Sarnıç Bağtepe, neck 74), remain undated. However,the first two of these are located near contemporaneouspalaeo-courses of the Gediz, and so are expected to morelikely cap individual Gediz terraces rather than wholestaircases. Westaway et al. (2004) and Maddy et al.(2005) assumed that the Sarnıç Bağtepe neck was alsolocated close to the contemporaneous course of theGediz. However, from the characteristic northwestwardflow direction of the basalt that erupted from it, alreadynoted, possibility now exists that it lay up to ∼1.5 km tothe south of this contemporaneous course, which (in thelight of earlier discussion) can be projected east-west

Notes to Table 3:Hco and Hmo are the initial thicknesses of the crust and mantlelithosphere. Tmo is the initial nominal Moho temperature. to is the starttime (before present) for increased erosion. Uo and U are the erosionrates before and after to. Uo is set to very small nonzero values to avoid‘division by zero’ runtime errors. Le, Ls, and Lh are the lengths, parallelto the sediment transport, of the eroding sediment source region, thedepocentre, and the ‘hinge zone’ in between. zwo and zw are theoffshore water depth at to and at present. Ed and Em are the assumedextensional strain rates for distributed deformation in the upper crustand mantle lithosphere. ηe is the predicted effective viscosity of thelower continental crust. Tm is the Moho temperature predicted fromeach fitted value of ηe using Westaway's (1998) preferred calibration.vu and Y are the predicted uplift rate of a marker that is not eroding, atpresent, and its predicted uplift since 3.1 Ma. For all models, densitiesof 1000, 2700, 3300, and 3100 kg m−3 are assumed for water, crust,mantle lithosphere, and asthenosphere, with 1.2 mm2 s−1 for κ, thethermal diffusivity of crust, and 9.81 m s−2 for g, the acceleration dueto gravity. In the absence of perturbations to the geotherm caused byerosion or sedimentation, zb for all model configurations would be12857.1 m; in the absence of corrections for subsequent crustalthickness changes, the values 3.1 million years after the start of allmodel runs would be zbe=12624.9 m and zbs=13092.1 m, makingΔzb=467.2 m.


between sites 26 and 4 in Fig. 7a. These two sites are∼2 km apart, so with an inferred downstream gradient of∼3 m km−1, comparable to the modern Gediz (cf.Westaway et al., 2004), a difference in height of ∼6 mbetween these sites can be expected, consistent with theavailable height control evidence. Earlier discussionraised the possibility that the basalt capping the easternSarnıç Plateau, Sürtmece (Fig. 11), originated fromBurgaz Bağtepe neck, not Sarnıç Bağtepe neck, and isthus most likely about two hundred thousand yearsyounger than the Sarnıç Bağtepe basalt. If so, the Gedizgravels beneath the eastern Sarnıç Plateau that are atlower levels than site 26 (i.e., those at sites 2, 3, 8, 17 and18; Fig. 7a) post-date the eruption of Sarnıç Bağtepeneck and instead record the early stages of the adjustmentprocess as the Gediz became diverted southward aroundits basalt, before this staircase of terraces itself became‘fossilised’ by the later Burgaz Bağtepe eruption thatcaused further southward diversion of the Gediz. Thisinterpretation may also explain why the palaeo-currentmeasurements at these sites in the eastern Sarnıç Plateauindicate southward palaeo-flow (Fig. 7a), not the west-ward direction that would be expected if the Maddy et al.(2005a) terrace correlation scheme were correct. Testingthis suggestion will require further dating and, in theabsence of more detailed height information, furtherdiscussion of how to improve this terrace correlationscheme is beyond the scope of this study.

As recognised by Westaway et al. (2004) and Maddyet al. (2005), the main value of the other three necks(İbrahimağa Bağtepe, neck 57; Toytepe, neck 73; andBurgaz Bağtepe, neck 75) is that by erupting at signif-icant distances from the contemporaneous valley floors,their basalts flowed over landscapes that had alreadybeen progressively incised by the Gediz and itstributaries, capping any river terraces that were present.However, since these basalts were unable to flow uphillfrom their sources, any terraces on the uphill side of eachof these necks did not become ‘fossilised’ by the basaltand have thus since been obliterated by erosion. Thefragmentary evidence of even higher Gediz terraces,reported by Westaway et al. (2003, 2004) farther ups-tream around Eynehan (Fig. 1), makes it clear thatterraces were formed at earlier times in this river system,even though none have yet been identified in the Kulaarea. Allowing for the difference in total incision be-tween Eynehan and Kula (∼360 against ∼410 m), thelowest of these very high terraces, ∼225 m above theGediz (Westaway et al., 2004), would correspond aroundKula to a relative level of∼260–250m. Inspection of themodel uplift histories in Fig. 12 suggests that such aterrace height would represent the incision since∼1900–

1800 ka, tentatively suggesting that this terrace formedin the span of time encompassing MIS 68 to 64. This agespan, covering the later part of the European Tiglianstage, is marked by terraces inmany rivers in western andcentral Europe (e.g., Westaway, 2002a); for instance, it isthe probable time of formation of the Stoke Row terraceof the Thames in southeast England (e.g., Whiteman andRose, 1992; Rose et al., 1999; Westaway et al., 2002).However, the localised and fragmentary character of thedeposits attributed by Westaway et al. (2003, 2004) tothese very high terraces will make it difficult to map theircounterparts systematically throughout theGediz system.

4.1. The role of Milankovitch forcing

It is well-established that during the Early Pleistocenethe global climate was dominated by ∼40 ka cyclicity,caused by fluctuations in the obliquity of the Earth'srotation axis relative to its orbital plane due to the grav-itational pull of the moon (e.g., Ruddiman et al., 1989;Raymo et al., 1989; deMenocal et al., 1993; Tiedemannet al., 1994). Marine oxygen isotope records from theeastern Mediterranean region (Fig. 15) also showoverwhelming predominance of ∼40 ka cyclicity atthis time (e.g., Kroon et al., 1998). However, such re-cords differ from those from the global oceans in severalrespects. First, the observed fluctuations in oxygen iso-tope ratios are much stronger in the Mediterranean Seathan in the low latitude oceans (Fig. 15). This has longbeen interpreted as a consequence of fluctuations indischarge into the Mediterranean basin of river waterdepleted in 18O (e.g., Rohling and Hilgen, 1991). Sec-ond, the Mediterranean records from this time showmuch more significant fluctuations at the ∼20 ka, orprecessional, periodicity than do those from the low-latitude oceans (Fig. 15). Third, stacked MediterraneanQuaternary marine sedimentary sequences are inter-spersed with sapropel beds (Fig. 15), which result fromreductions in salinity caused by influxes of freshwater. Inmany palaeoclimate reconstructions, these sapropelshave been regarded as consequences of variations indischarge by the River Nile linked to fluctuations in theintensity of the Indian Ocean monsoon, and thus inferredto mark warm stages (with the strongest evaporationfrom the low-latitude ocean) (e.g., Rossignol-Strick,1985). However, relatively little is known about theEarly Pleistocene hydrology of the Nile (cf. Said, 1981),making this point of view difficult to test. A furtherdifficulty with this view, evident in Fig. 15, is that manyof the sapropels do not coincide with warm stages of theoceanic oxygen isotope record. The sequence ofsapropels begins around 3.1 Ma (3040 ka according to

Fig. 15. Thick line shows oxygen isotope and sapropel chronology forODP site 967, south of Cyprus, adapted from Fig. 3 of Kroon et al.(1998). These oxygen isotope variations are compared with those forODP site 849 (Mix et al., 1995), the thin line, from the easternequatorial Pacific. Solid symbols indicate eastern Mediterraneansapropels; diagonal shading indicates sapropel ‘ghosts’ (i.e., bedswith trace evidence indicating obliteration of primary sapropels bydiagenesis). ⁎symbols indicate complete 40 ka stages of the marineoxygen isotope timescale in the Early Pleistocene that contain nosapropels. See text for discussion.


Hilgen, 1991; 3151 ka according to Kroon et al., 1998),suggesting a link to high-latitude climate change instead(e.g., Westaway, 2001, 2002a). These sapropels maythus relate to increases in discharge caused by melting oflocal ice sheets or permafrost in upland areas such as theAlps or the Caucasus (or, indeed, the Pontide or Tauridemountain ranges in Turkey), or possibly meltwater re-leased from the margins of Scandinavian lowland ice-sheets that have advanced southward across westernRussia and Ukraine into the catchments of Black Searivers such as the Dnieper and Don.

It is evident from Fig. 15 that most of the even-numbered cold stages in the part of the marine oxygenisotope record that encompasses the Gediz high terracestaircase correspond with at least one Mediterraneansapropel, the exceptions being MIS 52, 50 and 38. Thissequence is similar to that deduced independently by usfrom the high staircase of Gediz terraces. An alternativeapproach could thus be to ‘tune’ the Gediz high terracestaircase to the Mediterranean oxygen isotope/sapropelrecord, which would suggest, counting backwards intime, that terraces I–V mark MIS 28–36, terraces VI–Xmark MIS 40–48, and terrace XI marks MIS 54.However, a related problem is that one does not knowwhen in each climate cycle each of the Gediz highterraces aggraded (see Maddy et al., 2005, for discussionof this point). Nonetheless, the apparent correlationbetween Gediz river terraces and Mediterranean sapro-pels suggests the possibility of a cause-and-effectrelationship, such that both may be sedimentaryconsequences of the same series of climate instabilities.The youngest sapropel is well-dated to 9–6 ka, the EarlyHolocene climatic optimum, during which pollen evi-dence indicates pervasive deciduous vegetation in theeastern Mediterranean region, with the estimated annualrainfall up to ∼1300 mm (Rossignol-Strick, 1999),roughly double the typical present-day value for westernTurkey. This abundant vegetation would expected toreduce the sediment influx into the Gediz, and the as-sociated high discharge would thus be expected to causethis river to incise, potentially explaining the ∼5 m ofHolocene incision that is typical along this river withinthe Alaşehir Graben (Hakyemez et al., 1999); thuscreating accommodation space for subsequent sedimen-tation. This event may also explain the ∼5 m of incisionby the Gediz downstream of Kula Bridge (Fig. 5c),below the level of its youngest terrace, which pre-datesthe β4 basalt eruption, also the incision since theaggradation of the ∼5 m terrace at Palankaya (Fig. 3d).The deposits forming this low terrace may themselvespossibly date from the Younger Dryas stage (∼12 ka),when the easternMediterranean region was cold and arid

(e.g., Rossignol-Strick, 1999), leading to sparse vegeta-tion and thus an increased influx of sediment into anyriver system. If these phase relationships for the mostrecent climate cycle are applicable to the EarlyPleistocene, then for instance the incision to below thelevel of Gediz high terrace XI may have coincided withoffshore deposition of sapropel 38, near the end of MIS54 (Fig. 15), and the subsequent aggradation of fluvialgravel, forming this terrace, occurred sometime duringthe tens of thousands of years following this incisionevent, before the next incision that accompanied


sapropel 37 (MIS 49) or 36 (MIS 48) and created theaccommodation space for the gravel of Gediz highterrace X.

Alternatively, one can speculate that some of theearlier Mediterranean sapropels may mark relativelybrief cold periods caused by climate instabilities anal-ogous to the Late Pleistocene Heinrich events, sapropelspossibly representing melting events at the ends of thesephases. If so, then these events would be expected toaccompany reductions in vegetation cover, leading tosignificant influxes of sediment into rivers such as theGediz. The resulting high ratio of sediment flux todischarge would be conducive to terrace aggradation.Such an association with brief cold stages is consistentwith the known ages of climate instabilities revealed bylate Middle Pleistocene and Late Pleistocene river ter-races (Macklin et al., 2002) and stacked marine sedi-mentary sequences (Martrat et al., 2004) elsewhere inthe Mediterranean region, and may also account for thelarge numbers of Middle Pleistocene terraces observedin some river systems in central and NW Europe(Westaway, 2003). Given the Mediterranean sapropelrecord, the implication of such an association would bethat climate instabilities of this type have beenoccurring since ∼3.1 Ma. Such a relationship must beregarded for the time being as unproven, but ifconfirmed it would provide a natural explanation forwhy the global fluvial record indicates a systematicincrease in erosion rates around 3.1 Ma, linked toclimate deterioration (e.g., Westaway, 2002a; Bridglandand Westaway, in press), whereas evidence of ice-rafting indicates that the earliest large lowland ice sheetsadjoining the North Atlantic Ocean date from ∼2.75 Ma(e.g., Maslin et al., 1996). Further resolution of anycharacteristic timing of the formation of the high Gedizterraces in relation to Milankovitch forcing will requiremore detailed investigation, beyond the scope of thepresent study.

4.2. Implications of this modelling

Notwithstanding the approximations that have beenmade, as stated above and in previous publications (e.g.,Westaway, 2002c; Westaway et al., 2002), the modellingthat has been carried out provides a natural explanationfor the history of uplift and fluvial incision in thisregion. Climate deterioration around 3.1 Ma is inferredto have caused the start of erosion of older stacked LateCenozoic terrestrial sequences, this material being trans-ported westward by the Gediz and redeposited in part inthe Alaşehir Graben and in part in the Aegean Sea. Thecombined effect of this erosion and deposition was to

establish an eastward pressure gradient in the lower crust,driving mobile lower crust from beneath the AegeanSea to be beneath western Turkey, causing crustal thin-ning and subsidence in the former region and crustalthickening and surface uplift in the latter. The modellingassumes a spatial average erosion rate of 0.1 mm a−1

throughout, but in reality the erosion rate seems to haveincreased somewhat (to an estimated spatial average rateof ∼0.12 mm a−1) around the end of the Early Pleis-tocene, possibly as a result of the more severe coldstages that developed with the onset of ∼100 kaMilankovitch periodicity, although limitations of thetechnique used prevent this change in erosion rates frombeing modelled. This coupled process can thus accountfor the principal development of relief in this regionsince the Pliocene: the observed onshore uplift and off-shore subsidence, without any contribution from otherprocesses such as crustal extension.

One significant inference from this modelling is thatthe effective viscosity of the lower continental crust, ηe(as defined by Westaway, 1998), is low, being2×1018 Pa s for our preferred solution (M3 inFig. 14). As already noted (Table 3), using theWestaway (1998) calibration, this implies a Mohoviscosity of ∼3×1016 Pa s and a Moho temperature of∼660 °C in western Turkey. For comparison, geodeticstudies of post-seismic deformation following theAugust 1999 İzmit earthquake in northern Turkey(e.g., Hearn et al., 2002) indicate a Moho viscosity of∼1017 Pa s, suggesting a slightly lower Mohotemperature, consistent with the lower surface heatflow evident in northern Turkey relative to westernTurkey (cf. Tezcan, 1979; Ilkışık, 1995; Pfister andRybach, 1996; Pfister et al., 1998). The higher Mohoviscosity of ∼2×1017 Pa s deduced for western Turkeyin the rheological modelling by Westaway et al. (2004)makes less sense given the expected regional variationsin crustal properties, and is superseded by the presentresults. The difference in these estimated values ofMoho viscosity and ηe for western Turkey results in partfrom changes to the dataset that has been modelled andin part from changes (already noted) to the computerimplementation of the technique. The fact that the newset of results makes better overall regional sense than theold set suggests that the additional age and heightcontrols are beneficial and also that the improvements tothe technique are on the right track.

One significant difference between the results obtain-ed relates to the tendency of the former variant of thetechnique to predict that uplift rates of a sediment sourceregion that is eroding at a constant rate increase mono-tonically over time, unless prevented by the supply of


sediment running out. In the new variant uplift rates (andrates of subsidence of depocentres) initially increasefollowing a change in rates of surface processes, thenprogressively decrease as the system begins to recovertowards a new steady state. Inspection of Fig. 14 indi-cates that the time taken before this recovery beginsincreases with ηe; for instance, the time lag between theincrease in erosion rates and the fastest uplift responsevaries from 2725 ka (M1, ηe=5× 1018 Pa s) to 2100 ka(M2, ηe=3×10

18 Pa s), 1625 ka (M3, ηe=2×1018 Pa s)

and 1025 ka (M4, ηe=1× 1018 Pa s). This is because, putcrudely, the lower ηe is, the faster the lower crust flowsout from beneath the model depocentre, so the faster thecrust thins, reducing the excess pressure at the base of thebrittle layer, thus bringing the system closer to a steadystate. Alternative solutions (not shown) with a smallerlength scale thus adjust faster towards the new steadystate, because a given rate of erosion sets up a givenpressure differential between sediment source anddepocentre, but the closer they are together, the greaterthe corresponding pressure gradient, and thus the fasterthe lower-crust flows between these regions. Compari-son of Figs 13 and 14 indicates that imposing asubmarine rather than subaerial depocentre, keeping allother model parameters the same, means that the systemtakes longer to begin to adjust towards its new steadystate. This is because the water loading acts to put thesystem farther away from a steady state, as previouslynoted (Westaway, 2002c), so it takes longer to recover.The non-water-loaded solutions in Fig. 13 havedeveloped such that, by the present day, in each thelateral variation in zb,Δzb, is proportional to ηe, such thatthe ratio ηe /Δzb is almost constant. Given that a linearrheology is being assumed for the lower crust (seeWestaway, 1998) and that for these solutions the pressuredifferences between sediment source and depocentredepend only on Δzb, the observation (Fig. 13d) that themean horizontal velocity of the induced lower-crustalflow is very similar for all these solutions can beexplained. It is thus evident that these systems haveevolved into a delicate state of near-equilibrium forwhich rates of lower-crustal flow, and of surface upliftand subsidence, are virtually independent of the assumedrheology of the lower crust. In contrast, for the water-loaded solutions (Fig. 14), the situation is more complex,as the pressure differences between sediment sources anddepocentres depend both on Δzb and on the water loads.However, these solutions have also evolved such thatpredicted present-day process rates are almost indepen-dent of the assumed lower-crustal rheology.

These modelling results thus have some bearing onrecent debate in the geomorphological literature as to

whether landscape response to surface processes is or isnot typically in a steady state (e.g., Pazzaglia andKnuepfer, 2001; Bull, 2002). A casual examination ofthe Kula dataset might lead to the conclusion that ‘thetime-averaged incision rate by the Gediz has remained thesame, ∼0.15 mm a−1, during and since the Early Pleisto-cene and that this system is thus in a steady state’.However, setting aside the results of the above numericalmodelling, which make it obvious that it is not in a steadystate, one can also readily reach this conclusion directlyfrom the observational evidence, because the 0.15 mma−1 incision rate significantly exceeds the upper bound tothe spatial average erosion rate derived from the observedmaximum of ∼400 m of erosion since ∼3 Ma. Nonethe-less, if there was no evidence of the starting conditions(i.e., if the Ulubey Formation limestone had been erodedfrom everywhere in the region, not just from most of it)one would have no idea about what the starting conditionswere, and the question whether this landscape is in asteady state would thus be much more difficult to judge.

Rather than attempting to solve the physics under-lying the Late Cenozoic development of topography, ashas been the aim of this study and its predecessors (e.g.,Westaway, 2002c; Westaway et al., 2004), most of thequantitative geomorphological literature to date hasinstead adopted an empirical approach, for instancetrying to solve coupled systems of empirical equationsrepresenting factors such as conditions for slopestability and longitudinal equilibrium of river channels(e.g., Whipple, 2001; Willett et al., 2001). As Westaway(2004b) noted, no attempt has been made in such studiesto incorporate the physics of isostatic compensationunder non-steady-state conditions, which seems essen-tial for any rigorous solution of this problem. Further-more, most of this literature has investigated regions ofextreme erosion rates, such as Taiwan, southern NewZealand and Cascadia, the US Pacific Northwest. AsWestaway (2004b) also noted, such high erosion ratespreclude the preservation of evidence relating to theconditions at the start of the rapid erosion. Despite theirnaivety regarding consideration of the physics ofisostasy, such empirical studies have nonetheless suc-cessfully predicted many aspects of the development oftopography. Furthermore, it is now evident that scale-dependent effects identified in the present study mimicquite closely similar effects identified in previous em-pirical studies, such as that by Willett et al. (2001); forinstance, the smaller the spatial scale of an erodingsediment source region, the faster its topography adjuststowards a steady state after an increase in erosion rates.The key to understanding why such empirical studieshave been so successful may now be evident: as


illustrated in Fig. 13, if a landscape system is perturbedby an increase in erosion rates, and then the system isleft for long enough to recover towards its new steadystate, the recovery occurs in a manner that is virtuallyindependent of the assumed crustal rheology. The earlypart of this response will depend strongly on thisassumed rheology (as in Fig. 13), but if one has noevidence from that time (because it has all been eroded)one may find no basis for deducing any dependence onrheology, thus favouring an empirical approach instead.

5. Conclusions

A set of 13 new unspiked K–Ar dates has beenobtained for the Quaternary basaltic volcanism in theKula area of western Turkey, allowing improved datingof the fluvial deposits of the Gediz River that underliethese basalt flows. This dating is able, for the first time,to resolve different ages for the oldest basalts, assignedto category β2, that cap the earliest Gediz deposits rec-ognised in this area, at altitudes of ∼140 m to ∼210 mabove present river level. In particular, the β2 basaltcapping the Sarnıç Plateau (Fig. 2) is dated to 1215±16 ka (±2σ), suggesting that the youngest underlyingfluvial deposits,∼185m above present river level, are noyounger than MIS 38. In contrast, the β2 basalt cappingthe adjacent Burgaz Plateau is dated to 1014±23 ka,suggesting that the youngest underlying fluvial deposits,∼140m above present river level, date fromMIS 28. Thestaircase of 11 high Gediz terraces capping the latterplateau (Fig. 7a) is thus dated to MIS 48-28, assumingthey represent consecutive ∼40 ka Milankovitch cycles,as previously suggested (Maddy et al., 2005), although itis possible that as many as two cycles are missing fromthis sequence, such that the highest terrace is corre-spondingly older. Basalt flows assigned to the β3 cate-gory adjacent to these plateaus, capping Gediz terraces∼35 and ∼25 m above the present river level, have beendated to 236±6 and 180±5 ka, indicating incision ratesof ∼0.15 mm a−1, similar to the time-averaged ratessince the eruptions of the β2 basalts. The youngestbasalts, assigned to category β4, are Late Holocene; ourK–Ar results for them range from zero age to a maxi-mum of 7±2 ka.

This fluvial incision is interpreted using numericalmodelling as a consequence of uplift caused by aregional-scale increase in spatial average erosion rates to∼0.1 mm a−1, starting at ∼3100 ka, caused by climatedeterioration, since when a total of ∼410 m of uplift hasoccurred. Parameters deduced on this basis from theobserved disposition of the Early Pleistocene Gedizterraces include the local effective viscosity of the lower

crust, which is ∼2×1018 Pa s, the Moho temperature of∼660 °C, and the depth of the base of the brittle uppercrust, which is ∼13 km. The thin lithosphere in this arearesults in high heat flow, causing this relatively shallowbase of the brittle upper crust and the associatedrelatively thick lower-crustal layer, situated betweendepths of ∼13 and ∼30 km. It estimated that around900 ka, at the start of the ∼100 ka Milankovitch forcing,the spatial average erosion rate increased slightly, to∼0.12 mm a−1; the associated relatively sluggishvariations in uplift rates are as expected given therelatively thick lower-crustal layer.

This modelling indicates that the growth of topog-raphy since the Pliocene in this study region has notinvolved a steady state. The landscape was significantlyperturbed by the Middle Pliocene increase in erosionrates, as has subsequently adjusted towards—but notreached—a new steady state consistent with theseincreased erosion rates. It would not be possible toconstrain what has been occurring from the Middle toLate Pleistocene or even the Early Pleistocene upliftresponse alone; information regarding the startingconditions is also essential, this being available in thisregion from the older geological record of stackedfluvial and lacustrine deposition. This result has majorimplications for the rigorous modelling of uplifthistories in regions of rapid erosion, where preservationof information to constrain the starting conditions isunlikely.

Acknowledgments

This work contributes to International GeoscienceProgrammes 449 ‘Global correlation of Late Cenozoicfluvial deposits’ and 518 ‘Fluvial sequences as evidencefor landscape and climatic evolution in the LateCenozoic’. We thank the three anonymous reviewersfor their thoughtful and constructive comments.

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Late Cenozoic uplift of western Turkey: Improved dating and numerical modelling of the Gediz river...

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