Chronology of Neogene and Quaternary uplift and magmatism in the Caucasus: constraints from K–Ar dating of volcanism in Armenia

ELSEVIER

Tectonophysics 304 (1999) 157–186

Chronology of Neogene and Quaternary uplift and magmatism in the Caucasus: constraints from K–Ar dating of volcanism in Armenia John Mitchell a , Rob Westaway b,* a

Department of Physics, University of Newcastle upon Tyne, Newcastle upon Tyne, NE1 7RU, UK b 16 Neville Square, Durham, DH1 3PY, UK Received 25 February 1998; accepted 21 December 1998

Abstract The Greater Caucasus is one of Earth’s highest actively-uplifting mountain ranges; the adjoining Caspian Sea basin contains a substantial proportion of its hydrocarbon reserves. Like other parts of the former Soviet Union, the Neogene and Quaternary chronology of these important regions has not previously been well-defined. It has thus been impossible to obtain reliable estimates for rates of processes such as uplift of the Caucasus and sedimentation in the Caspian Sea. Previous studies have established the relative timings of events in the region, using correlation schemes between volcanism, glaciations, and the stratigraphy of the Caspian basin. However, a range of absolute chronologies has previously been proposed for these sediments and igneous rocks, based mainly on different interpretations of their magnetostratigraphic records. By K–Ar dating, we determine the ages of volcanism at three localities in Armenia as 1.1, 0.8 and 0.8 Ma. Using these data and other evidence, we propose a revision to the chronology of this region, in which a distinctive brief interval of normal magnetic polarity in the local sedimentary and volcanic magnetostratigraphic records is matched to the Cobb Mountain event in the global record rather than the Olduvai event or an earlier subchron as had previously been thought. We thus interpret a ¾1.5 Ma timing for the start of volcanism in the Lesser Caucasus, and also suggest a ¾1.2 Ma timing for the Late Akchagyl transgression of the Caspian Sea, a key event in the regional stratigraphy when this water body reached its greatest extent. We tentatively correlate this transgression with the melting event following glaciation during stage 36 of the oxygen isotope timescale, which was thus the first time during the Pleistocene when eastern Europe was covered by a lowland ice sheet. Time-averaged since ¾1 Ma, the flanks of the eastern Greater Caucasus mountains are shown to have uplifted at ¾0.6 mm a 1 and the Lesser Caucasus at ¾0.3 mm a 1 . We show that the rate and spatial scale of this uplift are too great to be the result of plate convergence, and suggest instead that this uplift is caused by crustal thickening due to inward lower-crustal flow to beneath these mountain ranges. At the start of magmatism in both the Greater and Lesser Caucasus, the estimated crustal thickness was ¾45 km. We thus suggest that this magmatism has been caused by heating of the mantle lithosphere due to earlier crustal thickening, the temperature rise required to initiate magmatism being the same in both cases.  1999 Elsevier Science B.V. All rights reserved. Keywords: K=Ar; Armenia; Caucasus; Quaternary; volcanism; vertical movements

Ł Corresponding

author. E-mail: [email protected]

0040-1951/99/$ – see front matter  1999 Elsevier Science B.V. All rights reserved. PII: S 0 0 4 0 - 1 9 5 1 ( 9 9 ) 0 0 0 2 7 - X

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1. Introduction The Caucasus region (Fig. 1) has experienced dramatic environmental changes over the past few million years. In the eastern Greater Caucasus mountains, fossil clifflines are present at ¾3.5 km altitude, with marine sediments of Late Sarmatian (Late Miocene) age preserved on the adjacent wavecut platform (e.g. Budagov, 1964). Most of the ¾4 km relief of this mountain range has thus developed in the past few million years in a region where only lowlands existed beforehand (Fig. 2). Existing reviews of this region in the western literature (e.g., Khain, 1975; Adamia et al., 1977; Zonenshain and Le Pichon, 1986; Philip et al., 1989; Zonenshain et al., 1990) are insufficiently detailed to provide starting points for any detailed study, making reference to the Russian language literature indispensable. Philip et al. (1989), for instance, proposed that uplift of the Caucasus began in latest Miocene time following a continental collision after the closure of an ocean which used to link the Black and Caspian Seas (Fig. 1). Zonenshain et al. (1990)

and Cloetingh and Burov (1996) summarised similar interpretations. However, this hypothesis, that the uplift of these mountain ranges is the isostatic response to crustal shortening which accommodates convergence between the Arabian and Eurasian plates, can only be tested if uplift rates are known. Many studies have been carried out in the Russian language literature to try to explain the topography of the Caucasus and other elevated regions within the former Soviet Union, particularly by modelling gravity data (e.g., Artem’yev and Balavadze, 1973; Artem’yev et al., 1985). For instance, Artyushkov (1973, 1974) suggested several possible explanations, including the presence of anomalously low-density upper mantle, possibly due to dynamic upwelling and=or active chemical differentiation at greater depths. On the other hand, Artem’yev et al. (1985) argued that gravity modelling requires anomalously high-density upper mantle beneath the Caucasus. Of course, the inference of mantle density anomalies through gravity modelling depends on the density distribution assumed in the continental crust. Like Philip et al. (1989), we see no reason to

Fig. 1. Location map of our study region (star), also showing localities outside this region which are discussed in the text. Fine dotted lines denote international frontiers. Large arrows mark overflow channels which carried glacial meltwater into and out of the Caspian region, which is now isolated. The area covered by Fig. 3 is outlined. Following the collapse of the Soviet Union, many place names have been changed. Changes relevant to this and later maps, as well as to place names mentioned in the text, are listed in Table 1.

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Table 1 Placename conversions

Fig. 2. Map showing the palaeogeography of the Caucasus region in Middle Sarmatian (i.e., late Middle Miocene) time. This map is non-palinspastic: no attempt is made to restore subsequent plate motions or continental deformation. Bold line denotes the palaeocoastline, redrawn from Nalivkin (1973, fig. 213), with ornament on the landward side. Note that at this time, the highest land in the Greater Caucasus formed an island within the eastern part of the Paratethys Sea, which was itself isolated from the global marine environment. Similar maps have also been published by others (e.g., Steininger and Ro¨gl, 1984).

question the alternative interpretation that the region consists of continental crust of normal density, up to ¾60 km thick, overlying mantle of normal density. The initial aim of this study is thus to constrain the timing of some key volcanic and sedimentary events in this region, and to thereby estimate uplift rates. We constrain the timing of the Quaternary volcanism in Armenia, and work through correlation schemes between this volcanism, glaciations, and highstands of the Caspian Sea, to investigate the timing of the other processes. Having established a chronology for the region, we then investigate possible explanations of the observed uplift. One key result concerns the Late Akchagyl highstand of the Caspian Sea, when this water body reached its greatest extent (e.g., Moskvitin, 1961) and distinctive marine sediment was deposited over much of European Russia and central Asia. Soviet studies have regarded this event as Late Pliocene, around 2.5 Ma, or at ¾1.8 Ma near the Plio– Pleistocene boundary (see below). On the contrary, some western correlation schemes regard it as Early Pliocene (e.g., Steininger and Ro¨gl, 1984), considering it equivalent to the transgression of the Mediterranean Sea following its Messinian lowstand. As is

Modern name

Older name

Aragats (volcano) Gyandja Vanadzor Gyumri Mtkvari (river) Nalband Tsovagyugh Sevan (town) Vladikavkaz Zemo-Karabulakhi

Alago¨z Kirovabad Kirovakan Leninakan Kura Shirakamut Tsovak Yelenovka Ordzhonikidze Zurtaketi

As the Armenian alphabet has 38 letters, Russian has 33 and English 26, many spelling variants occur for instance depending on whether one transcribes from Armenian to English direct or via Russian. These are omitted.

explained in detail below, our dating evidence and the existing magnetostratigraphic data set together indicate instead that this Akchagyl highstand occurred at ¾1.2 Ma. This effectively compresses the young part of the Caucasian stratigraphy in time by a factor of ¾2, indicating uplift about twice as fast as was previously thought. We realise that the original dating evidence presented in this study is limited. We indeed regard this study as the first step in establishing a robust quantitative chronology for this region by isotopic dating involving more detailed sampling of its late Neogene and Quaternary igneous rocks over a wider region. It may indeed astonish readers that our small number of samples can contribute to dating such an important region. The reasons why this is possible are summarised here, and are explained in more detail below. First, this region has been isolated from the global marine environment for many millions of years. As a result, it is not possible to use micropalaeontology to match its stratigraphy against global standards. Second, because of this isolation, its Neogene and Quaternary chronology was originally established as a series of local stratigraphic definitions, which were arbitrarily defined — without supporting calibration evidence — as equivalent to divisions of the global stratigraphy. Third, the available isotopic and fission track dates from this region are discordant: the scatter in apparent ages for individual sites is much greater than the quoted er-

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Table 2 Available K–Ar dates for Plio–Quaternary magmatism in the Greater Caucasus

1 2 3 4 5 6 7

Rock

Mineral

[K2 O] a

[40 Ar* ] b

Granophyre Granophyre Granodiorite Granodiorite Granophyre Granophyre Granophyre

Biotite Biotite Biotite Biotite Biotite Biotite Biotite Anorthoclase Biotite Whole rock Whole rock Whole rock Sanidine Biotite Sanidine Sanidine Whole rock

7.87 8.35 8.27 8.04 8.89 8.47 8.10 4.94 8.63 4.37 3.10 2.19 9.44 8.37 6.93 8.36 3.21

1:092 ð 10 1:098 ð 10 0:717 ð 10 0:666 ð 10 0:633 ð 10 0:538 ð 10 0:459 ð 10 0:297 ð 10 0:437 ð 10 0:230 ð 10 0:084 ð 10 ? 0:907 ð 10 0:795 ð 10 0:616 ð 10 0:655 ð 10 ?

8 9 10 11 12

Granophyre Dacite Dacite Granophyre Grey Rhyolite

13 14 15

Black Rhyolite Rhyolite Dacite

3 3 3 3 3 3 3 3 3 3 3 3 3 3 3

Cc

Age (Ma)

Locality

52–60 53–58 78–80 82–85 65–85 80–85 60–75 80–90 80–86 85–90 95–96 ? 45–70 70–75 70–80 84–86 ?

4:30 š 0:20 4:07 š 0:20 2:69 š 0:20 2:57 š 0:20 2:21 š 0:20 1:97 š 0:15 1:76 š 0:15 1:86 š 0:15 1:57 š 0:15 1:63 š 0:15 0:84 š 0:25 0.69 2:98 š 0:20 2:94 š 0:20 2:75 š 0:20 2:43 š 0:20 0.50

Mt. Kairobi, near Rioni R. valley (W. Georgia) Mt. Tsurungal, near Chenishkali R. valley (W. Georgia) Midagravin Glacier (North Osetia, Russia) Dzhungusu R. valley (trib. of Chegem R., Balkaria, Russia) Mt. Tepli (North Osetia, Russia) Kyrtyk R. valley (trib. of Baksan R.), Balkaria, Russia Eljurta granite (Baksan R. valley), Balkaria, Russia Same place Mt. Kal’ko (Khevsurskoy Aragvi R. valley), E. Georgia Dike piercing 7 Stock piercing 5 Mt. Taymazi, near Tanadon R. valley Chegem R. valley, Balkaria, Russia Same place Flow covering 12 Same locality as 4 Mt. Elbrus (western border of Balkaria, Russia)

Data are from Arakelyants et al. (1968). To facilitate comparison, we have converted their values into the units used for our own dating, have recalculated the ages using modern decay constants from Steiger and Ja¨ger (1977), and have provided additional location information. a [K O] denotes the potassium oxide concentration (in wt%). 2 b [40 Ar* ] denotes the concentration of radiogenic 40 Ar (in mm3 g 1 ). c C denotes the percentage of 40 Ar that is attributable to atmospheric contamination.

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ror bounds, indicating the presence of some form of systematic error; and in some cases noted previously (e.g., by Khaburzaniya et al., 1979) the sequence of existing K–Ar dates is not in stratigraphic order. The discordance in the available isotopic dates from this region means that they cannot be used to define a unique chronology, as others (e.g., Khaburzaniya et al., 1979) have already noted. Previous commentaries (e.g., Rast, 1971) have identified problems affecting the isotopic dating literature of the former Soviet Union, due to a combination of technological limitations and difficulties with quality control. Because part of the point of the present study was to try to understand the cause of the discordance in the available dates from the Soviet literature for the young volcanism of the Lesser Caucasus, we chose the same method (i.e. whole rock, K–Ar dating), but using modern analysis procedures. A further difficulty arises, because K–Ar dating articles in the Soviet literature often presented results without technical details of the method or the data used to calculate the results. The reader is instead referred to separate reports, produced by dating laboratories, for this information. Following the collapse of the Soviet Union and the closure of many of its scientific establishments, these reports are now inaccessible. In the text below we have indicated where published dates cannot now be verified because the supporting technical documentation is lacking. In the absence of such documentation, one could choose to regard all such dates as void. However, we mention these dates for two reasons: because their range indicates the extent of discordance in the existing set of dates; and because they have been used to interpret results and define previous chronologies. In a further attempt to assess the reliability of Soviet-era K–Ar dates, we list in Table 2 examples from the Greater Caucasus, which we have recalculated using modern decay constants and expressed using the same units and format as our own results. As is discussed in more detail below, this sample of Soviet-era K– Ar dates which are concordant with other evidence are based on quantities of radiogenic argon at least ten times greater than are measured in our present study. Samples such as ours, whose low potassium contents and young ages lead to minimal radiogenic argon, appear instead to have been beyond the dating capability of Soviet-era facilities.

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The technical difficulties in using the K–Ar method to date young basalts (containing, say, ¾1% of potassium and with ages of, say, ¾1 Ma) have led some former Soviet laboratories to try other methods. For instance, fission track dating has been used to determine the time since basalts have cooled after erupting subaerially (e.g., Ushko et al., 1987). In the western literature this method is normally used instead to date cooling due to exhumation of rocks. However, many Soviet fission track dating studies also fall short of the minimum standards of documentation listed by Hurford and Green (1982), due to the lack of explanation of details of the experimental method and procedure for age calculation.

2. Background information As previous studies (e.g., Westaway, 1990) have noted, uplift rates in the Caucasus are poorly constrained, because different data yield conflicting estimates. Evidence of river valley incision by hundreds of metres since the last glaciation has been used to infer uplift rates of 10 mm a 1 or more (e.g., Rastvorova and Shcherbakova, 1967). However, there is no simple relationship between rates of incision and uplift: calculating the isostatic response to valley incision requires detailed analysis of the topography (e.g., Gilchrist et al., 1994) which has not yet been attempted in the Caucasus. Geodetic evidence of vertical motion at ¾10 mm a 1 or more is also suspect, as many benchmarks have been sited on unconsolidated ground (e.g., Lilienberg, 1967). Sediments deposited during transgressions of the Caspian Sea are found higher within the Greater Caucasus than in other coastal localities (e.g., Lilienberg, 1967). However, it is difficult to use this information to estimate uplift rates, because the ages of transgressions are disputed (see below). Because the Caspian region has been isolated from the global marine environment since Middle Miocene (Serravallian) time (e.g. Van Couvering and Miller, 1971), a detailed local stratigraphy has been devised (see, e.g., Nalivkin, 1973). We only consider its main divisions: Sarmatian, Meotian, Pontian, Kimmerian, Akchagyl, Apsheron, and Baku. During Miocene time, the Greater Caucasus formed an island within the landlocked Paratethys Sea, which

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linked Bavaria and the modern Aral Sea region (Fig. 2). By late Middle Miocene (Early Sarmatian) time, no external connection to the Paratethys Sea remained (e.g., Van Couvering and Miller, 1971), creating an isolated sea with reduced salinity (e.g., Nalivkin, 1973, p. 633). In Late Sarmatian time uplift eliminated the marine connection between the Black and Caspian Seas south of the Greater Caucasus (Nalivkin, 1973, p. 630). Continuing uplift of the Earth’s surface and=or drawdown of water level north of the Greater Caucasus in the Meotian and Pontian gradually restricted the remaining linkage between these seas (Nalivkin, 1973, p. 632). Correlation of these locally-defined stages of time with the global stratigraphy is difficult. Although the base Pliocene has been defined in the Soviet literature as the base Pontian, the similarity in fossil content of many Pontian and Meotian sediments suggests that no major boundary occurs at this point and the base Pliocene is earlier (e.g., Nalivkin, 1973, p. 630). However, in western correlations such as Steininger and Papp (1979), the Sarmatian spans 13.3–10.7 Ma, the Meotian 10.7–9.6 Ma, the Pontian 9.6–5.2 Ma, the Kimmerian 5.2–3.3 Ma and the Akchagyl 3.3–1.8 Ma, thus placing both the Pontian and the Meotian within the Miocene. The Early Sarmatian is tied to the global microfossil stratigraphy and absolute timescale at ¾14–13 Ma (Van Couvering and Miller, 1971) or 13.3 Ma (Steininger and Papp, 1979). The simplest argument for such a date uses andesitic lavas and tuffs in western Ukraine and eastern Slovakia, K–Ar dated at ¾13 Ma, which are interbedded with marine sediments containing Early Sarmatian fossils (e.g., Mikhaylova et al., 1974; Steininger et al., 1976). Furthermore, the pattern of geomagnetic polarity reversals in the Early Sarmatian (e.g., Mikhaylova et al., 1974) fits the sequence during chron 5A, which spans ¾13–12 Ma in modern timescales such as Berggren et al. (1985). In the Kimmerian, regressions occurred in the isolated Caspian and Black Sea basins (e.g., Karlov, 1961; Kvasov, 1964). The Caspian Sea surface regressed to ¾ 500 m, leaving a hypersaline lake in its southern part (Kvasov, 1964) with delta of the palaeo-Volga river near Baku (Fig. 1). The up to 4 km of deltaic sand and silt which accumulated, forms the main hydrocarbon reservoir in Azerbaijan, the Productive Suite (e.g., Khain, 1950; Nalivkin,

1973, pp. 630–631). During the later Akchagyl transgression, the Caspian Sea flooded north up the valleys of the Volga and its tributaries almost to the Ural mountains (e.g., Gorelov, 1962) and eastward to the Aral Sea and beyond up the Amu river almost to the Afghan border and up the Syr to the vicinity of Tashkent (e.g., Zhivotovskaya and Popov, 1969). Its outlet to the Black Sea was via the Manych valley (e.g. Maslyaev, 1979) (Figs. 1 and 3). Elsewhere, its shoreline is traced at ¾180–200 m altitude in the Volga–Urals region (e.g., Gorelov, 1962), near sea-level around Kara-Bogaz gulf (e.g., Kleyner, 1965), and in the subsurface more than 500 m below sea-level in western Turkmenistan (e.g., Rastsvetayev, 1969). However, at localities such as Derbent and Shemakha in the eastern Greater Caucasus (Fig. 3), Akchagyl marine sediments reach ¾850–900 m altitude (e.g., Lilienberg, 1967; Nesmeyanov and Voeikova, 1994). These differences of many hundreds of metres indicate the magnitude of the responses since Akchagyl time to all the processes which have affected the isostatic balance, including crustal thickness changes, denudation, sedimentation, and ice loading in adjacent areas. In 1956 the Baku was defined as the earliest Pleistocene stage and the Apsheron the latest Pliocene (Nalivkin, 1973, p. 623). However, this local definition bears no relation to the international stratotype for the Plio–Pleistocene boundary using marine sediments in southern Italy (e.g., Aguirre and Pasini, 1985). Evidence of both Akchagyl and earlier mountain glaciations exists within the Greater Caucasus (e.g., Nalivkin, 1973, pp. 630–631, 635). However, the Akchagyl highstand of the Caspian Sea was associated with the earliest lowland glaciation of eastern European Russia (e.g., Moskvitin, 1961). Evidence for this, such as ice-rafted debris and erosional trace evidence of drift ice along the Akchagyl palaeocoastline, comes notably from near Kazan (e.g., Moskvitin, 1961) (Fig. 1). Greenland was of course forested around 1.8 Ma and did not become ice-covered until later (e.g. Funder et al., 1985). It is thus unlikely that eastern Europe carried a lowland ice sheet at ¾1.8 Ma or earlier, raising the possibility that established chronologies for the Akchagyl are in error. Following Akchagyl time, the Caspian Sea level has fluctuated. Important transgressions during the

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Apsheron and Baku stages mark extremes of river discharge into the Caspian Sea during the melting of icecaps in European Russia and western Siberia (e.g., Moskvitin, 1961). During the peak of the most recent Khvalyn transgression at ¾15–12 ka, the Caspian Sea reached 47 m above global mean sea level (75 m above its present level) allowing it to overflow into the Black Sea via the Manych palaeo-strait (e.g., Kaplin et al., 1972; Gerasimov, 1978; Grosswald, 1980; Fig. 3). At this time the Caspian Sea was fed not only as at present by rivers draining eastern Europe, but also from western Siberia via the Turgai overflow channel (Fig. 1) and from central Asia (e.g., Grosswald, 1980). It is evident from the above paragraphs that the chronologies of glaciations and sea-level variations in the Caspian Sea are linked: the melting of continental icesheets has caused runoff into the Caspian Sea via the Volga and rivers in central Asia. Moskvitin (1961) correlated the Akchagyl with the Oka glaciation, which was then thought to be the earliest in European Russia but which is now assigned to stage 12 of the oxygen isotope timescale around 0.45 Ma (e.g., Velichko and Faustova, 1986). Although older glaciations are now known (e.g., Velichko and Faustova, 1986), the complexity of the evidence has so far precluded any overall glaciation scheme, let alone one incorporating river terraces and the Caspian Sea.

3. Volcanism of the study region The young volcanism of the Caucasus has been discussed even less in the western literature than its stratigraphy. Reviews by Khain (1975), Zonenshain and Le Pichon (1986), Philip et al. (1989), and others indicate its spatial extent and report a Neogene to Quaternary timing with more or less continuous activity. This view reflects early Soviet studies (e.g., Paffengol’ts, 1959; Milanovskiy and Koronovskiy, 1961; Nalivkin, 1973) which were based in part on intuitive arguments using field relationships. More recent isotopic dating studies indicate, on the contrary, that magmatism has been concentrated at a particular time in each locality (see below). This young magmatism has involved the formation of basalt flows, stratovolcanoes, and hypabyssal

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intrusions which have become unroofed by rapid erosion (Fig. 3). The stratovolcanoes form a northward prolongation of a chain of similar edifices in eastern Turkey (e.g., Su¨phan and Tendu¨rek in Fig. 3). Data from Milanovskiy and Koronovskiy (1961), Blumenthal et al. (1964), and Nalivkin (1973) establish that Aragats has erupted mainly basalts, basaltic andesites, and dacites, whereas Elbrus has produced mainly dacites and rhyolites. They thus continue the northward increase in silica content across eastern Turkey noted by Pearce et al. (1990). Some studies (e.g., Zonenshain and Le Pichon, 1986; Philip et al., 1989; Zonenshain et al., 1990) have suggested that this volcanism has been caused by subduction of an ocean basin which formerly linked the Black and Caspian Seas between the Caucasus ranges. Its distribution thus requires two subducting slabs, dipping north beneath the Greater Caucasus and south beneath the Lesser Caucasus. On the contrary, Platt and England (1993) have proposed that magmatism in mountain belts is evidence of delamination of the lower part of the thickened mantle lithosphere, which they estimated will initially thicken during crustal shortening before becoming unstable due to its high density, and the resulting heating of what remains of the lithosphere. Others have proposed different causes: for instance, Yılmaz et al. (1987) suggested that the young volcanism in eastern Turkey, adjacent to the Lesser Caucasus, has resulted from the heating of the lower continental crust or mantle lithosphere which accompanies crustal thickening. Bogatikov et al. (1992) and Lipman et al. (1993) have noted geochemical similarities between the Pliocene magmatism of the Greater Caucasus and the Early Tertiary magmatism of the western USA which has been widely regarded as a consequence of earlier crustal thickening. Koronovskiy and Demina (1996) suggested, using geochemical arguments, that heating caused by crustal thickening may also explain the young volcanism of the Lesser Caucasus. 3.1. Magmatism of the Greater Caucasus and its surroundings The Greater Caucasus range is formed of Palaeozoic granite intruded into older crystalline basement (e.g., Shevchenko, 1972). Younger magmatism, in-

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Fig. 3. Map of the Caucasus region, showing features mentioned in the text. See Fig. 1 for location. The area covered by Fig. 4 is outlined. Asterisk symbols mark hypabyssal igneous intrusions. Triangle ornament marks the Akchagyl shoreline of the Caspian Sea in the vicinity of the Manych outflow channel, from Maslyaev (1979). Volcanoes and other mountains are distinguished by different ornament. Numbers 1 to 6, circled, mark localities in southern Georgia where Quaternary basalts have been magnetostratigraphically dated. Localities 1 and 2 mark the gorges of the Karabulkh and Mashavera rivers near Zemo–Karabulakhi and Zemo–Orozmani. Locality 3 is Akhalkalaki, 4 is Kumurdo, 5 is Khertvisi, and 6 is the Godzerdi Pass.

cluding the Chegem extrusives, the Eljurta granite, and the associated Tyrnyauz ore mineralization, was K–Ar dated by Arakelyants et al. (1968) (Table 2). It has been re-examined using the 40 Ar=39 Ar and Rb– Sr methods (Zhuravlev and Negrey, 1993; Gazis et al., 1995), whose results are concordant with this K– Ar dating. This magmatism was concentrated around 2.8 Ma, with ages of 2.0–1.9 Ma marking the cooling of intrusive rocks below the closure temperature for argon. In contrast, the young rhyolitic=trachytic magmatism of the Pyatigorsk area, farther north (Fig. 3), has been K–Ar dated at 9–8 Ma (Borsuk et al., 1989). Young volcanism persists for ¾250 km along the

western Greater Caucasus from Elbrus (altitude 5642 m) to Kazbeg (5047 m) (Fig. 3). Kazbeg is considered no longer active, but some eruptions of Elbrus post-date the last glaciation and fumaroles near its summit emit a vapour plume (e.g., Masurenkov and Panteleyev, 1962; Blumenthal et al., 1964; Nalivkin, 1973, p. 668–670). 3.2. Volcanism of the Lesser Caucasus and its surroundings Aragats volcano (Figs. 2 and 3) occupies a ¾1000 km2 area of western Armenia, rising from ¾1500 to 4095 m. It is considered inactive (Blumenthal et al.,

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1964), and its summit carries an icecap (e.g., Louis, 1943), although minor eruptions occurred early in historical time (Nalivkin, 1973, p. 672). Ararat, 90 km south of Aragats, is a double-peaked stratovolcano with summits at 5165 and 3908 m, also with an area of ¾1000 km2 . Like Aragats, it is covered by an icecap (Birman, 1968) and is considered inactive (Blumenthal et al., 1964). An early basalt was K–Ar dated by Pearce et al. (1990) at 1:51 š 0:19 Ma, whereas one of its youngest cones yielded a K–Ar age of ¾0.5 Ma (Sanver, 1968). Ararat and Aragats (Fig. 4) cover much of the Ararat sedimentary basin, which contains up to ¾4 km of Mesozoic and Tertiary sediment overlying metamorphic basement (Burshtar and Tolmachevskiy, 1965). East and north of Aragats, other volcanoes form the Gegham plateau (highest summit, Azhdahak, 3597 m; Fig. 4) and the Javakhet plateau (highest summit, Abul, 3304 m; Fig. 3). Magnetostratigraphic work has been carried out in both localities (e.g., Bol’shakov and Solodovnikov, 1969, 1980; Vekua et al., 1977; Khaburzaniya et al., 1979), plus limited isotopic dating of Javakhet (e.g., Vekua et al., 1977; Khaburzaniya et al., 1979). Tuffs which have erupted from these volcanic edifices cover much of the surrounding landscape. In addition, many river valleys in the Lesser Caucasus are partly filled by lavas which erupted locally. An example is the Akera (Fig. 3), where conglomerate, sand and silt are interbedded with tuffs and basaltic and andesitic lavas (Pashady and Suleymanov, 1973). According to Pashady and Suleymanov (1973), these sediments can be correlated with fossiliferous beds elsewhere in Azerbaijan using their content of characteristic minerals, enabling a chronology to be established despite the absence of absolute age evidence. From the proportions of volcanogenic material, they concluded that volcanism began in the Akchagyl stage but became most intense around the Apsheron–Baku boundary. A stratigraphy for the young volcanism of Armenia was established by Paffengol’ts (1959). This involved five suites, A to E, which were thought to be separated by the four Alpine glaciations (Gu¨nz, Mindel, Riss and Wu¨rm) in the classical Penck and Bru¨ckner (1909) chronology. As the oldest lavas are directly underlain by lacustrine sediments, with no intervening glacial deposits (Nalivkin, 1973, pp. 629, 672–675),

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the earliest volcanism evidently pre-dates the earliest local glaciation. However, as this original Penck and Bru¨ckner (1909) scheme was never properly defined and contained internal inconsistencies even within the Alps, and many more Alpine glaciations are now known, it should be abandoned (e.g., Sibrava, 1986). Its use has nonetheless continued in Armenia: subsequent studies (e.g., of magnetostratigraphy) have taken the ages of lavas directly from their designations by Paffengol’ts (1959) (for instance, a suite C flow supposedly has ‘Mindel–Riss’ age). Bol’shakov and Solodovnikov (1969) showed that suite A flows at localities B and C in Fig. 4 are reverse-magnetised but all stratigraphically younger flows are normally-magnetised. Because the youngest transition from reverse to normal magnetic polarity (R ! N) had been reported in Caspian Sea sediments in the Late Apsheron stage (e.g., Mammedov, 1967), which was defined as part of the Pliocene, Bol’shakov and Solodovnikov (1969) regarded the suite A volcanics as Pliocene. It is via this type of correlation that reviews have widely reported the start of volcanism in Armenia as Pliocene, an example being the Middle Pliocene or ¾3.5 Ma timing quoted by Zonenshain and Le Pichon (1986).

4. Data analysis and interpretation 4.1. Data collection and analysis Samples for isotopic dating were collected (by R.W.) in February 1989. About 500 g of material was recovered at each site, from which ¾10 g specimens were prepared for analysis. Samples were collected near Yerevan, east of Aragats volcano near Oganavan, and west of Spitak near Nalband (Fig. 4). Specimens were analysed at the Department of Physics, University of Newcastle upon Tyne, using the procedure of Wilkinson et al. (1986) (Table 3). Concentrations of potassium were determined using an EEL 450 flame photometer with a lithium internal standard (Table 3). Concentrations of the argon isotopes 36 Ar, 38 Ar, and 40 Ar were determined using a modified Kratos MS10 mass spectrometer coupled to a UHV gas extraction line. Isotope dilution analyses, using a measured ‘spike’ of 38 Ar, enabled components of atmospheric and radiogenic argon to

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Table 3 Potassium–argon dating Locality

Sample

[K2 O] a

[40 Ar* ] b

Yerevan Oganavan Oganavan Nalband Nalband

1A1 1B1 2A1 3B2 3B3

1:91 š 0:02 2:34 š 0:01 2:44 š 0:01 3:26 š 0:03 3:09 š 0:03

6:89 š 0:50 ð 10 5:51 š 1:11 ð 10 6:62 š 0:53 ð 10 9:47 š 0:89 ð 10 7:51 š 0:56 ð 10

5 5 5 5 5

Cc

Age d (Ma)

91.1 97.7 91.9 91.0 88.9

1:12 š 0:08 0:73 š 0:15 0:84 š 0:07 0:90 š 0:08 0:75 š 0:06

Samples are identified by names allocated in the field. a [K O] denotes the concentration of potassium oxide (in wt%), as the mean of three analyses. 2 b [40 Ar* ] denotes the concentration of radiogenic 40 Ar (in mm3 g 1 ), as the mean of two analyses. c C denotes the percentage of 40 Ar that is attributable to atmospheric contamination, reported as the higher of two measured values. d Age calculations use Eq. 1. Note that the calculated age does not depend on C, which is quoted here simply to demonstrate that radiogenic argon is present in the specimens. The percentage error estimates quoted for radiogenic 40 Ar do depend on C, as discussed in the text, on the uncertainties in measurements of individual argon isotopes present (which are determined individually for each specimen, rather than using the nominal estimates quoted in the text), and on other uncertainties in the calibration of the apparatus used against internal and external standards.

be distinguished, using the procedure described by Dalrymple and Lanphere (1969, pp. 54–65). This method uses the fact that radiogenic argon is pure 40 Ar, whereas atmospheric argon is a mixture of 36 Ar, 38 Ar, and 40 Ar in known proportions. The calculated percentage uncertainty in radiogenic argon content depends on the percentage of contamination of the specimen by atmospheric argon. With the apparatus used, uncertainties in 38 Ar and 40 Ar measurements are typically ¾0.1%; the margin for the less abundant isotope 36 Ar being ¾0.4%. Standard calculations (e.g., Baksi et al., 1967) establish that with these margins of uncertainty, for a specimen with 90% atmospheric 40 Ar the direct contribution of uncertainties in individual argon isotope measurements to the calculated concentration of radiogenic argon is ¾4%. Ages t were determined using the K–Ar dating equation:

Ł ArŁ tD f ½e [K ] ð 40

(1)

from Dalrymple and Lanphere (1969, p. 49), where [K ] and [40 Ar* ] are the quantities of potassium and radiogenic 40 Ar in a specimen, ½e is the decay constant for the decay of 40 K by electron capture to 40 Ar, and f is the relative abundance of 40 K in potassium. Standard values of f and ½e , from Steiger and Ja¨ger (1977) are used. For the data in Table 3, the percentage errors in [K ] are much smaller in magnitude than the percentage errors in [40 Ar* ]. The resulting percentage error in age is thus very similar to the percentage error in [40 Ar* ], being typically ¾10% for our specimens. 4.1.1. Yerevan Sample 1A1 was collected beside the Yerevan– Ashtarak road, ¾5 km NW of Yerevan city centre

Fig. 4. Map of western Armenia (Hayastan) and surrounding parts of Georgia, Azerbaijan, Iran, and Turkey, showing our sampling sites, the major volcanic edifices of Aragats, Gegham and Ararat, and adjacent key localities discussed in the text. See Fig. 1 or Fig. 3 for location. Letters A to D denote localities mentioned in the text. A is the Avan borehole near Yerevan. B and C mark Kamo and Tsovagyugh, sites of magnetostratigraphic studies by Bol’shakov and Solodovnikov (1969). D marks Karbi, where the basalt flow which we dated at Oganavan had previously been studied by Bol’shakov and Solodovnikov (1969). Dotted line indicates the route followed between sampling sites. Active faults are not shown, as detailed maps and field photographs showing these in the area of our sampling sites have already been published by Westaway (1990) and others. Many of these faults control the drainage geometry of the region. The Leninakan fault, mentioned in the present text, strikes east–west and is situated ¾5 km north of Gyumri (former Leninakan). Compiled from sources cited in the text, with most information from figs. 231 and 233 of Nalivkin (1973). Geological detail is omitted for the part of this map area which is outside the former Soviet Union.

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Fig. 5. Schematic cross-section through Yerevan showing the relationship which we interpret between the basalt exposed at site 1A1, in the Hrazdan gorge, and in the Avan borehole. Borehole evidence is from Nalivkin (1973, pp. 625, 629).

and just outside its built up area. Leaving the city centre, this road descends into the Hrazdan river gorge, incised into basalt. The sides of this ¾30 m deep gorge are cut into basalt, except for the uppermost ¾5 m which are in alluvium. The sampling site is ¾1.5 km west of this gorge, within the flat interior of the Ararat sedimentary basin. At Avan (A in Fig. 4), ¾10 km east of this site, a borehole penetrated 60 m of alluvium then 61 m of basalt, which is underlain by sediments of the Ararat basin but no older basalt (Nalivkin, 1973, pp. 625, 629). The field and borehole evidence suggests the interpretation in Fig. 5, with the same sequence of flows at Avan, in the Hrazdan gorge, and at our sampling site. As the borehole evidence at Avan and the exposure in the Hrazdan gorge indicate no sedimentation between flows, we deduce that magmatism in the Yerevan area was concentrated in time. However, as our sample was obtained from a point late in the sequence of flows, the start of local magmatism was evidently somewhat earlier. The sample is of a vesicular basalt with ¾1 mm phenocrysts of olivine, calcic plagioclase, and augite. The matrix is partly glassy, containing many small bubbles. The vesicles are also partly lined with glass, some of which has devitrified. We estimate the proportion of devitrified material within the specimen as a few percent. 4.1.2. Oganavan Sample 1B1 was collected north of Oganavan near the northern edge of the Ashtarak lava flow. Sample 2A1 was collected about 300 m farther south within

the same outcrop. This flow emerges from the eastern flank of Aragats volcano, then turns gradually south on entering the Kasakh river gorge, reaching the town of Ashtarak. It thus partly backfilled this pre-existing gorge, but has since been eroded to form a newer gorge bounded by vertical cliffs. This flow was assigned to suite D by Paffengol’ts (1959), a supposed ‘Riss–Wu¨rm’ age. Magnetostratigraphic work by Bol’shakov and Solodovnikov (1969) at Karbi (D in Fig. 4) indicates that it has normal polarity. Samples 1B1 and 2A1 are similar in thin section, comprising vesicular basalt with phenocrysts of plagioclase, magnesium-rich olivine, augite, and hypersthene, with a glassy matrix containing small crystals of plagioclase, augite and olivine. Farther west, higher up the volcano near the village of Aragats, this basalt is overlain by dacitic flows and tuffs (Fig. 4), which also have normal magnetic polarity according to Bol’shakov and Solodovnikov (1980). 4.1.3. Nalband Welded tuff samples 3B2 and 3B3 were collected ¾1 km east of Nalband on the eastern side of the gorge of the Chichkhan river, which flows south into the Pambak just east of the reverse fault that slipped in the 1988 Spitak earthquake (see Westaway, 1990, Fig. 7). This site is thus in the hanging-wall of this fault, where local uplift requires the rivers to incise their courses. This tuff crops out over a wide area between these young gorges (e.g., Westaway, 1990). West of Nalband in the footwall of this fault the Pambak is instead ponded, creating an alluvial plain. Both samples contain fragments and whole crystals of plagioclase, augite, hypersthene, and rare olivine, plus clasts of basalt and rhyolite or dacite. The glassy matrix is partly vesicular and partly formed of individual fragments which shows signs of flow banding and deformation. Similar welded tuffs of the Tumanyan suite from farther north (Fig. 4), described by Gushchin (1994), have varied compositions including some resembling our samples. 4.2. Interpretation Sample 1A1 from Yerevan was dated to 1:12 š 0:08 Ma (Table 3). This basalt thus formed during the Matuyama chron between the Jaramillo and Olduvai events.

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Samples 1B1 and 2A1 from the Ashtarak flow at Oganavan gave dates which are concordant at the one standard deviation (1σ) level, of 0:73 š 0:15 Ma and 0:84 š 0:07 Ma. As this flow is normally magnetised (Bol’shakov and Solodovnikov, 1969), they thus could in principle be either early Brunhes or late Jaramillo. We prefer the Brunhes alternative, first, because the normal polarity of the overlying dacitic flow is incompatible with a late Jaramillo age for the Ashtarak flow, as one would then expect the younger flow to be reversed. Second, the 1σ error bounds formally rule out a Jaramillo age but are concordant with the time scale of Shackleton et al. (1990) where the Brunhes starts at 0.78 Ma. Tuff samples 3B2 and 3B3 from Nalband gave dates of 0:90 š 0:08 Ma and 0:75 š 0:06 Ma. These are concordant at 2σ and yield a weighted mean of 0:81 š 0:11 Ma which is indistinguishable from the age of the basalt at Oganavan. We have carried out several tests on the validity of our reported dates. First, our dates from the same flows have been checked for inter-sample concordance, which is generally regarded as the best reliability criterion (e.g., Dalrymple and Lanphere, 1969). Second, inspection of our samples in hand specimen and thin section revealed no evidence of weathering effects sufficient to potentially indicate that a closed system for argon has not been maintained. Third, in each locality the basalt or tuff has remained at the Earth’s surface since formation: it has neither been buried nor reheated by younger magmatism. These samples have thus remained exposed to the atmosphere. In contrast, had they been buried below the local water table the likelihood of weathering would be greater. Armenia is arid, with annual rainfall of only 304 mm at Yerevan (Lydolph, 1977, p. 426). Exposed basalt and tuff samples are thus not expected to weather significantly. 4.3. Discussion Although all our specimens are very fresh-looking in hand specimen and thin section, we noted some evidence of devitrification in one of them (1A1; see above). Devitrification may of course be caused by weathering and may lead to argon loss. However, it is not necessarily a weathering process — or even a water-dependent process — and can be deuteric (e.g.,

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Barker, 1983). Furthermore, devitrification does not necessarily lead to argon loss: instances are known where samples containing high proportions of devitrified glass yield potassium–argon dates which do not differ significantly from those of non-devitrified specimens from the same volcanic suite (e.g., Baksi et al., 1967). Nonetheless, other instances are known where devitrified specimens do yield discordant ages (e.g., Schaeffer et al., 1961). Thin section evidence is anyway not an infallible criterion for whole rock dating of basalts: it has long been established that “the most reasonable method of assessing the reliability of the age of a lava is to date samples from various parts of the body and examine the consistency” (Baksi et al., 1967). Samples from two of our sites pass this test for consistency. For the third site, we are confident that any systematic error in age due to devitrification is small compared with the random uncertainty quoted in Table 3, given our estimate that only a few percent of sample 1A1 has devitrified. Due to their low potassium contents and young ages, our samples contain little radiogenic argon (Table 3), less than 10% of the amounts in most of the Plio–Quaternary rocks in the Greater Caucasus that have previously been successfully dated using the K–Ar method (Table 2). This small set of isotopic dates does however indicate that there is no technical obstacle to systematically dating the Quaternary magmatism of the Lesser Caucasus using the K–Ar method. 4.4. Deductions from our K–Ar dating Our dating at Oganavan indicates that the ‘suite D’ basaltic magmatism of Aragats, its youngest large-scale activity, occurred at ¾0.8 Ma, the same age as for the tuffs which we have dated at Nalband. Although the available historical record (e.g., Nalivkin, 1973, p. 672) indicates that minor volcanism has persisted into the Holocene, the present volcanic landscape in these localities had thus largely developed by ¾0.8 Ma. This timing contrasts with the ¾9–8 Ma activity near Pyatigorsk and the ¾3 Ma peak activity in the western Greater Caucasus, mentioned earlier. As suites B and C also have normal magnetic polarity (Bol’shakov and Solodovnikov, 1969, 1980), it appears that the bulk of the activity of Aragats — comprising suites B to D — erupted

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over a brief interval of time around the start of the Brunhes chron. At Yerevan, we have dated to ¾1.1 Ma one of the uppermost of a sequence of basalt flows within the Ararat sedimentary basin, which appear from the absence of sedimentary interbedding (Fig. 5) to have erupted over a relatively short interval of time. The start of magmatism in this locality may match the ¾1.5 Ma age of the early basalt from Ararat volcano, ¾40 km farther south, which was K–Ar dated by Pearce et al. (1990). As Westaway (1990) noted, the distances by which Pleistocene welded tuffs are offset can indicate the vertical slip rates on reverse faults in Armenia. Westaway (1990) estimated the slip rates on several such faults assuming a nominal age for this tuff of 1.5–1.0 Ma. As its age is now measured as ¾0.8 Ma, improved estimates for these slip rates can now be made. For instance, Westaway (1990) described the Leninakan fault, a north-dipping reverse fault situated north of Gyumri (former Leninakan), across which the tuff is offset by 250 m. Westaway (1990; Fig. 7) provided a detailed location map showing this and other active reverse faults in the region. Assuming the age of this tuff to be 1.5–1.0 Ma, Westaway (1990) estimated a vertical slip rate of 0.17–0.25 mm a 1 , whereas the revised age of ¾0.8 Ma indicates 0.31 mm a 1 . The conclusions of Westaway (1990) about the order-of-magnitude of slip rates on reverse faults in Armenia and their minor role in taking up Arabia–Eurasia plate convergence thus remain largely unaffected.

5. Chronology of the Caucasus–Caspian region 5.1. Previous dating schemes Previous dating schemes for the sediments of the Caspian basin have been based largely on magnetostratigraphy. One of the earliest schemes, by Mammedov (1967), placed the youngest R ! N transition (i.e., transition from reverse to normal geomagnetic polarity) at the Middle–Late Apsheron boundary and the previous N ! R transition in the Early–Middle Akchagyl, with many older transitions in the preceding sequence of redbeds which were deposited in the Kimmerian around the re-

stricted Caspian basin. If his youngest transitions are regarded as Gauss–Matuyama and Matuyama– Brunhes, then the Akchagyl stage spans ¾3–2 Ma. Later work (e.g., Gurariy et al., 1977) recognised that the Middle–Late Apsheron R ! N transition was followed by a N ! R transition, with a further R ! N transition near the Late Apsheron– Baku stage boundary. The brief normal-polarity event in the Middle–Late Apsheron was identified as the Jaramillo, placing the Matuyama–Brunhes boundary near the Late Apsheron–Baku stage boundary. Gurariy et al. (1977) accepted the Middle Akchagyl N ! R transition as the Gauss–Matuyama, and a preceding Base Akchagyl R ! N transition as the Gilbert– Gauss. They interpreted a brief Early Akchagyl normal-polarity event, evident at some sites, as the Kaena and=or Mammoth event. They also noted evidence of two brief normal-polarity events around the Middle– Late Akchagyl and Akchagyl–Apsheron boundaries, which they interpreted as the Re´union and Olduvai events. The Akchagyl–Apsheron boundary is thus dated to the Olduvai subchron (e.g., Gurariy et al., 1977) such that the Akchagyl stage spans ¾3.6 to 1.8 Ma (Fig. 6). Other studies (e.g., Steininger and Papp, 1979; Ushko et al., 1987) also correlate the Late Akchagyl with ¾1.8 Ma. Fission track dating by Ushko et al. (1987) justified this interpretation with a Late Akchagyl date of 2:0 š 0:3 Ma and Early Apsheron dates of 1:7 š 0:2 and 1:8 š 0:1 Ma, these dates being cooling ages of tuffs interbedded in Akchagyl sediments. However, their article falls short of the minimum standards of documentation listed by Hurford and Green (1982), due to the lack of explanation of the experimental method and procedure for age calculation. Steininger and Papp (1979) justified their Akchagyl chronology using K–Ar dates of 2.26 Ma and 0:95 š 0:3 Ma, one of which is too young to fit their scheme. Furthermore, they regarded the Kimmerian as a time of highstand within the Caspian Sea, ignoring work (e.g., by Kvasov, 1964) which demonstrates that it was a time of lowstand. Their correlation scheme evidently breaks down during the Pliocene. 5.2. Proposed revision It is evident that the previous dating schemes have been based on matching parts of the sedimentary se-

J. Mitchell, R. Westaway / Tectonophysics 304 (1999) 157–186

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Fig. 6. Comparison of the existing interpretation scheme for the Caspian region from the Soviet literature with our proposed alternative. The second column shows the magnetostratigraphy of Caspian sediments, from Gurariy et al. (1977). Black and white shading denote normal and reversed magnetic polarity; grey indicates uncertain polarity. The first and third columns interpret these data in terms of subchrons and absolute ages, using the time scale calibrations by Shackleton et al. (1990) and Hilgen (1991a,b). The first column indicates the established correlation scheme (e.g., Gurariy et al., 1977); the third indicates our proposed alternative. The fourth column summarises our interpretation of the Quaternary volcanism of the Lesser Caucasus, which is consistent with columns two and three, showing our interpreted timings of volcanism at key localities discussed in the text. The concentration of volcanism around the Matuyama–Brunhes boundary, when (for instance) the bulk of flows forming Aragats volcano erupted, is indicated schematically by thicker ornament. See text for discussion.

quence around the Caspian basin to the global record of magnetic polarity reversals, working backwards in time. Past omission of any subchrons will thus

cause systematically old magnetostratigraphic ages for earlier parts of the sedimentary record. Although the more recent dating schemes recog-

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Table 4 Proposed chronology Event

Oxygen isotope time scale correlative

Age (Ma)

Top of Baku stage End of Early Baku substage Start of Brunhes chron End of Apsheron stage Base of Emilian stage End of Jaramillo event Start of Jaramillo event Early Apsheron highstand Cobb Mountain event End of Akchagyl stage End of Olduvai event Start of Olduvai event

end of stage 12 end of stage 16 end of stage 20 end of stage 22 start of stage 25 mid stage 27 mid stage 31 end of stage 34 start of stage 35 end of stage 36 start of stage 63 start of stage 71

0.45 0.62 0.78 0.87 0.96 0.99 1.07 1.13 1.19 1.20 1.71 1.95

Baku, Apsheron, and Akchagyl chronologies are proposed by this study. Other ages are from Shackleton et al. (1990) or Hilgen (1991a). See text for discussion.

nise some normal-polarity subchrons within the Matuyama chron (Fig. 6), no scheme for the Caspian basin has incorporated the Cobb Mountain event at ¾1.2 Ma, which has been widely recognised elsewhere (e.g., Shackleton et al., 1990; Cande and Kent, 1995). One would not expect this event to be registered everywhere, but it can be expected within the Caspian basin where sedimentation has been continuous and rapid. The main feature of our proposed revision to the chronology is to correlate the brief interval of normal magnetic polarity around the Akchagyl– Apsheron stage boundary with this Cobb Mountain event, rather than — as has previously been thought — with the Olduvai event. This adjusts the age of this stage boundary from ¾1.8 Ma to ¾1.2 Ma. Fig. 6 and Table 4 show a revised chronology of the magnetostratigraphic record in which the known sequence of polarity reversals from Caspian sediments is mapped onto a complete sequence of subchrons. In this revised scheme, the base Akchagyl is around the start of the Olduvai event at ¾2 Ma. 5.3. Implications for the chronology of magmatism in the Lesser Caucasus The stratigraphic evidence indicates a Late Akchagyl start of volcanism in the Akera valley of SW Azerbaijan (Fig. 3) (Pashady and Suleymanov,

1973). Using our revised dating scheme, the start of volcanism in this locality is dated to ¾1.5 Ma. It thus roughly matches the K–Ar date for the start of volcanism for Ararat (Pearce et al., 1990). The chronology of the widespread young volcanism of southern Georgia has been investigated in detail by Vekua et al. (1977) and Khaburzaniya et al. (1979), using magnetostratigraphy supported by limited K–Ar and palaeontological dating. After sampling at many sites, and correlating individual lavas between sites, Vekua et al. (1977) established a composite magnetostratigraphic section involving five stages, with N ! R ! N ! R ! N polarities, and interpreted the start of magmatism as Early Pliocene. Khaburzaniya et al. (1979) reported an additional reverse-polarity stage at the start of the sequence, although this was identified at only one locality (locality 6 in Fig. 3), around the Godzerdi pass near the western end of the Lesser Caucasus, whose stratigraphy is difficult to correlate with other localities farther east. By matching the sequence of polarity reversals without including subchrons, they deduced an even earlier start of volcanism, around 10 Ma in the Late Miocene. Vekua et al. (1977) also deduced prolonged volcanism using geomorphological arguments. For instance, some sequences of lava flows (e.g., at Khertvisi, locality 5 in Fig. 3) accumulated without incision, then were incised by rivers, then partly backfilled by younger flows, then re-incised. Vekua et al. (1977) and others regarded the flows which predate the original incision as Pliocene and the younger flows in the modern gorges as Quaternary. However, at the measured incision rates of gorges in the Caucasus, in excess of 10 mm a 1 (e.g., Rastvorova and Shcherbakova, 1967), a 1 km deep gorge could form in only ¾0.1 million years. Lava filling such a gorge could thus readily fall within the same subchron as older lava in its surroundings. Attempts have also been made at palaeontological dating of these lavas. For instance, at some sites, such as Zemo-Karabulakhi and Zemo-Orozmani (localities 1 and 2 in Fig. 3), the earliest lavas cover lacustrine sediments containing mammal and plant remains of Akchagyl age (Khaburzaniya et al., 1979). Available K–Ar dating evidence for the young volcanism of the Lesser Caucasus in Georgia is extremely limited. Khaburzaniya et al. (1979) reported

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K–Ar dates of 0.3–0.5 Ma for the youngest lavas in this area, but provided no source for this information. At Akhalkalaki (locality 3 in Fig. 3), a suite of dolerite flows records a R ! N transition which has been interpreted as the Matuyama–Jaramillo (e.g., Khaburzaniya et al., 1979). However, at Kumurdo (locality 4 in Fig. 3) this dolerite is overlain by an andesitic flow of reverse polarity, for which a K–Ar date of 1.1–1.2 Ma from an inaccessible report was cited by Vekua et al. (1977). In these localities the sequence of interpreted ages is thus incompatible with the stratigraphic order. We suggest that the R ! N ! R ! N ! R ! N sequence of magnetic polarity reversals reported in these lavas is Matuyama–Cobb Mountain– Matuyama–Jaramillo–Matuyama–Brunhes, placing the start of magmatism in southern Georgia around ¾1.5 Ma. This is consistent with the Akchagyl age of the underlying lacustrine sediments given our revised chronology (Fig. 6), and indicates that magmatism began at roughly the same time throughout the Lesser Caucasus. 5.4. Implications for the chronology of glaciations Given that the Late Akchagyl highstand coincided with the earliest lowland glaciation of eastern European Russia (e.g. Moskvitin, 1961), it is now possible to estimate in which stage of the oxygen isotope time scale this glacial event occurred, as follows. Because the Cobb Mountain event falls within the transition from glacial stage 36 to interglacial stage 35 of the oxygen isotope time scale (Shackleton et al., 1990; Worm, 1997), this definition implies that the Late Akchagyl=Early Apsheron highstand of the Caspian Sea marked the melting event at the end of oxygen isotope stage 36, and thus that the earliest lowland glaciation of eastern European Russia occurred during stage 36. The Late Apsheron highstand of the Caspian Sea is required by both our chronology and the previously established chronology to fall at or near the end of the Matuyama chron after the Jaramillo event. It is thus most likely associated with the melting event at the end of oxygen isotope stage 22 that involved a particularly severe glaciation (e.g. Shackleton et al., 1990). The Early and Late Baku highstands of the Caspian Sea, in the Early Brunhes, now fit the

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melting events following oxygen isotope stages 16 and 12 which also marked severe glaciations (e.g., Shackleton et al., 1990). 5.5. Implications for regional uplift rates The evidence, from Budagov (1964), indicating Late Sarmatian age marine sediments at ¾3500 m altitudes, together with geomorphological features interpreted as fossil marine clifflines and offshore wavecut platforms, indicates a subsequent time-averaged uplift rate of ¾0.3 mm a 1 (Table 5). In contrast, the altitude limits of Quaternary marine sediments and Pleistocene and Holocene shorelines (Table 5) indicate much higher uplift rates, of ¾0.5 to 1.0 mm a 1 , even when generous allowances are made for the uncertainties in the contemporaneous palaeo-water-levels (see Table 5 caption). Furthermore, the Quaternary shorelines are all along the flanks of the eastern Greater Caucasus, whereas the Sarmatian site is on their crest (Fig. 3). The Kura basin to the south of the eastern Greater Caucasus, and the adjoining floor of the Caspian Sea, have formed depocentres throughout Quaternary time, where in some localities several kilometres of sediment have accumulated (e.g., Nalivkin, 1973). From gravity data, the thickness of continental crust tapers from ¾60 km along the axis of the eastern Greater Caucasus to ¾35 km beneath these depocentres (e.g., Philip et al., 1989). The ¾4 km altitude of the axis of the eastern Greater Caucasus is roughly as is expected for continental crust (density ¾2700 kg m 3 ) which has become thickened to ¾60 km thickness, in isostatic equilibrium above mantle lithosphere which overlies asthenosphere (density ¾3200 kg m 3 ), provided the mantle lithosphere has not itself thickened (e.g., Westaway, 1995). In these circumstances, the surface uplift rate is ¾1=6 of the crustal thickening rate in each locality (e.g., Westaway, 1995). The Quaternary sites on the flanks of the eastern Greater Caucasus are thus expected to be uplifting less rapidly than the range axis. Although no direct evidence exists to constrain them, Quaternary uplift rates along the axis of the eastern Greater Caucasus may well thus exceed 1 mm a 1 , indicating an even greater disparity relative to the post-Sarmatian time-averaged uplift rate (Fig. 7). Table 5 also lists two uplift estimates for the

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Table 5 Uplift rates Reference a

Feature

Locality

t (ka)

H (m)

Eastern Greater Caucasus 0, 1 Limit of Late Sarmatian marine sediment 2 Limit of Akchagyl marine sediment 2 Limit of Akchagyl marine sediment 3 Limit of Apsheron marine sediment 2 Limit of Apsheron marine sediment 3 Baku shoreline 3 Baku shoreline 4, 5 Early Khvalyn shoreline 4 Neocaspian shoreline

Shakhdag Derbent Izberbash Shemakha Izberbash Shemakha Makhachkala Baku Baku

10700 3500 1200 700 1200 900 870 800–850 870 550 450 300–400 450 240 15 62 3–4 19

Lesser Caucasus 1, 6 Limit of Meotian sediment 7 Limit of Akchagyl marine sediment

Hrazdan valley Akera valley

9600 1500 1200 500

Ho b (m)

U (m)

vu (mm a 1 )

0 š 100 3500 š 100 150 š 100 550 š 100 150 š 100 750 š 100 100 š 100 725 š 125 100 š 100 450 š 100 50 š 100 300 š 150 50 š 100 240 š 100 47 15 22 3

0:33 š 0:01 0:46 š 0:08 0:63 š 0:08 0:83 š 0:14 0:52 š 0:11 0:67 š 0:33 0:53 š 0:22 1:00 š 0:07 0:87 š 0:13

500 š 100 1500 š 100 150 š 100 350 š 100

0:21 š 0:01 0:29 š 0:08

Inferred ages t of Caspian palaeoshoreline features at altitude H are used to calculate time-averaged uplift rates vu as vu D U=t with U, the estimated uplift, calculated as H Ho where Ho is the assumed initial altitude. a References cited are: 0, Budagov (1964), 1, using Steininger and Papp (1979) chronology; 2, Nesmeyanov and Voeikova (1994); 3, Lilienberg (1967) or Dumitrashko and Lilienberg (1967); 4, Gerasimov (1978); 5, Kaplin et al. (1972); 6, Nalivkin (1973); 7, Pashady and Suleymanov (1973). Ages used for Sarmatian and Meotian shorelines mark the ends of these stages of geological time in the chronology of Steininger and Papp (1979). If the Meotian in fact falls within the Pliocene, as Nalivkin (1973, p. 630) tentatively suggested, then a greater subsequent time-averaged uplift rate is required. Ages for the Akchagyl, Apsheron, and Baku shorelines are from Table 4. Ages for the Early Khvalyn and Neocaspian shorelines are based on radiocarbon dating described in the references cited. A nominal uncertainty of š1 ka is assumed for the Early Khvalyn shoreline. b For the Neocaspian and Early Khvalyn highstands of the Caspian Sea, these shoreline levels are measured in localities which are distant from the Caucasus (e.g., Gerasimov, 1978). The Sarmatian shoreline level is assumed to have been the same as global sea level, because the Paratethys Sea had only recently become isolated from the global marine environment (e.g., Steininger and Ro¨gl, 1984). The Meotian level of 500 m is a nominal value based on the sedimentary evidence of a substantial drawdown in water level in the Caspian basin at this time (e.g., Nalivkin, 1973, pp. 630–632). The Akchagyl shoreline level of the Caspian Sea is assigned a nominal value of 150 m, close to the present-day 180–200 m altitude of this shoreline in the Volga–Urals region of Russia (Gorelov, 1962), which is assumed to not be uplifting significantly. The Apsheron and Baku shoreline levels are interpolated at nominal 50 m intervals between this Akchagyl shoreline level and present-day global sea level. Margins of uncertainty of š100 m are assigned to each of these nominal shoreline levels, and lead to the margins of uncertainty in the calculated uplift rates.

Lesser Caucasus. The first of these is based on the presence of Late Akchagyl age marine sediments in the Akera valley of SW Azerbaijan (Pashady and Suleymanov, 1973) at altitudes up to ¾500 m. Assuming our revised timing for the Late Akchagyl, a subsequent time-averaged uplift rate of ¾0.3 mm a 1 is indicated. A second uplift rate estimate can be made from the present-day altitudes of sediments in the Ararat basin of Armenia. The basalt in the Avan borehole near Yerevan (Fig. 5), at ¾1000 m present-day altitude, is underlain by 15 m of the Zangian formation of sand and clay with a Caspian-type mollusc fauna, then 400 m of evaporites and bituminous clay (e.g.

Nalivkin, 1973, pp. 625, 629). In the gorge of the Hrazdan river north of Yerevan, the Zangian instead reaches a thickness of ¾100 m (Nalivkin, 1973, pp. 625, 629) and is found at ¾1500 m altitude. Nalivkin (1973, pp. 624–629) reported conclusions of earlier work, which assigned the Zangian to the Meotian stage and the evaporites to the Middle– Upper Miocene and interpreted these sediments as having accumulated in an embayment linked to the Caspian basin of the Paratethys Sea. The estimate of a ¾0.2 mm a 1 time-averaged uplift rate for these sediments (Table 5) is based on this interpretation. However, Nalivkin (1973, pp. 624–629) also noted that Pliocene ages for these Zangian deposits were

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Fig. 7. Graph of uplift measurements for the eastern Greater Caucasus, showing the data in Table 5 and two predictions of the uplift variation using Eq. 3. Dashed line [1], fitted to the Quaternary data from the flanks of the eastern Greater Caucasus, is calculated assuming an initial surface altitude of 700 m, a start of uplift at 8.4 Ma, and K D 0:214 Ma 1 . Solid line [2] is an alternative solution for the uplift of the crest of the eastern Greater Caucasus, with a start of uplift at 6 Ma and K D 0:3 Ma 1 . See text for discussion.

also feasible, in which case the subsequent timeaveraged uplift rate has been greater. Despite this uncertainty, it is evident that the Quaternary uplift rate has been much lower in the Lesser Caucasus than in the Greater Caucasus.

6. Implications for partitioning of convergent plate motion Our results (Table 5) can be compared with uplift rates in more southerly parts of the zone of convergence between the Arabian and Eurasian plates. Such rates have been explained as a consequence of crustal thickening caused by plate convergence and thus used to estimate where this convergence is concentrated. In the Zagros mountains of Iran, the Earth’s surface was below sea-level until the Middle Miocene (Sto¨cklin, 1968) (or the earliest Late Miocene — at ¾12–11 Ma — according to Steininger and Ro¨gl, 1984). The ¾2 km present-day altitude of the Zagros

has thus developed at a time-averaged rate no greater than ¾0.2 mm a 1 . This evidence has been used in isostatic calculations (e.g. by Jackson and McKenzie, 1988) to support claims that the bulk of convergence between the Arabian and Eurasian plates is taken up by crustal thickening that balances shortening in the Zagros. Their argument is as follows. Standard theory for the isostatic response to crustal thickening (e.g., Westaway, 1995) requires the surface uplift rate to be no greater than one sixth of the crustal thickening rate. A crustal thickening rate of at least 1.2 mm a 1 is thus required to explain an uplift rate of 0.2 mm a 1 . Taking the north–south width of the zone of surface uplift in the Zagros mountains as ¾400 km, and the crustal thickness outside this zone as ¾40 km, a simple mass balance calculation requires crustal shortening at ¾1.2 mm a 1 ð400=40 or 12 mm a 1 . The convergence rate between the Arabian and Eurasian plates in the vicinity of the Caucasus and western Zagros mountains is ¾15 mm a 1 (e.g., Westaway, 1994a). Most of the plate convergence

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thus appears to be taken up by crustal shortening in the Zagros mountains. Even if one assumes an uplift rate in the Greater Caucasus no greater than ¾0.5 mm a 1 , a major problem arises if this uplift is to be explained by crustal thickening in response to convergence between the Arabian and Eurasian plates. Taking the north–south width of this zone of surface uplift as ¾200 km, mass balance requires crustal shortening across the Greater Caucasus at ¾15 mm a 1 , sufficient to take up all the relative motion between these plates. It thus follows that the uplift across the Zagros mountains and the Greater Caucasus, as well as across the areas in between, is too rapid across too wide an area to be a consequence of crustal thickening caused by plate convergence. Another process is evidently responsible for much of this uplift, and rates of uplift, whether in the Caucasus or the Zagros, provide no direct indication of the partitioning of plate convergence. Geodetic measurements using the Global Positioning System now also provide some constraint on shortening rates across the Caucasus. Reilinger et al. (1997) measured rates of northward motion, relative to a reference frame defined by stations elsewhere in the Eurasian plate, at six points in the present study region: Kars in northeastern Turkey (Fig. 3), Garni near Yerevan in Armenia (Fig. 4), three sites in Georgia between the Greater and Lesser Caucasus, and a site in southern Russia. Relative to their Eurasian reference frame, they deduced northward motion at 1 š 5 mm a 1 (š two standard deviations) at the Russian site, 2 š 5, 4 š 5, and 5 š 5 mm a 1 at the Georgian sites, 11 š 5 mm a 1 at Garni, and 7 š 3 mm a 1 at Kars. As there are no major active faults between Yerevan and Kars (Fig. 3), it is appropriate to average together their velocities to obtain 9 š 3 mm a 1 as the best estimate of the overall convergence rate between the southern margin of the Lesser Caucasus and the Eurasian reference frame. Reilinger et al. (1997) estimated that their reference frame may differ from the actual motion of the stable interior of the Eurasian plate, possibly by as much as several millimetres per year. These data thus indicate a shortening rate of several millimetres per year, at least, across the Lesser Caucasus, and a shortening rate across the Greater Caucasus which

is too low to resolve and may well be zero. Like the geological data, they indicate the need to explain crustal thickening in the Caucasus without crustal shortening.

7. Atectonic uplift of the eastern Greater Caucasus It is evident that the surface uplift in the Caucasus has been caused by crustal thickening, but this thickening has not been caused directly by plate motions. Following Westaway (1994b), we thus suggest an alternative possibility, that the inflow of lower crust beneath the eastern Greater Caucasus which has caused the surface uplift has been driven by the isostatic response to denudation of this mountain range and sedimentation in its surroundings. Crustal deformation of this sort, caused by lower-crustal flow, is called atectonic deformation (e.g., Kaufman and Royden, 1994), in contrast with tectonic deformation which is caused by plate motions. Although many factors, including climate and lithology, influence rates of denudation (i.e., mechanical erosion plus chemical weathering) of mountain ranges, Summerfield and Hulton (1994) and others have established that the most important overall factor is relief. Their results also indicate rough proportionality between denudation rates and the altitude of mountain ranges, which exert the main control on local relief. To determine the resulting predicted form of uplift variation, we note from Westaway (1994b) that the sediment flux leaving an eroding region is proportional to the denudation rate, and the inflow of lower crust to beneath an eroding region is proportional to the uplift rate of any marker horizon which has not eroded. Given the result of Summerfield and Hulton (1994) that the denudation rate is proportional to altitude, if the inflow of lower crust to beneath an eroding region is proportional to the sediment flux leaving it, then the rate of change of altitude of any marker horizon dS=dt is proportional to the altitude S of the mountain range, or: dS  vu D K S dt

(2)

where K is a scale factor. This equation has the

J. Mitchell, R. Westaway / Tectonophysics 304 (1999) 157–186

solution: S D So exp [K .t

to /]

(3)

where So is the initial surface altitude and to is the time of the start of uplift. From the data of Budagov (1964) indicating a 3500 m Sarmatian shoreline level in a locality where summits rise to 4200 m, So was 700 m and S is now 4200 m. From the uplift data for Holocene shorelines of the Caspian Sea (Table 5), vu is now 0.9 mm a 1 , so from Eq. 2 K is 0.214 Ma 1 , and from Eq. 3, t to is ¾8.4 Ma for the present day. Solution [1] in Fig. 7 shows the predicted uplift profile with these parameter values. Solution [2] shows an alternative uplift profile, in which uplift with K D 0:3 Ma 1 is assumed to have started at 6 Ma (see below). The predicted present-day uplift rate for this profile is 1.2 mm a 1 , a plausible estimate for the present-day uplift rate along the axis of the eastern Greater Caucasus given the uplift rate of 0.9 mm a 1 along the flanks (Table 5). In continental crust, the base of the upper-crustal brittle layer occurs at the depth where the temperature is ¾300ºC. Below this depth, the crustal material deforms plastically, and pressure gradients caused by lateral variations in the depth at the base of the brittle layer will determine its sense of flow. Westaway (1998) has shown that the flux Q of lower crustal material per unit horizontal distance perpendicular to the flow can be expressed as: QD

PW3 12e

(4)

where W is the thickness of the plastic lower-crustal channel, P is the horizontal pressure gradient driving the flow, and e is the effective viscosity of the lower continental crust. The local viscosity at each point in the lower crust will decrease with depth due to the geothermal gradient. Westaway (1998) showed that e can be estimated as ¾50 times the viscosity at the Moho, and is ¾1019 Pa s for crust of normal thickness and with a normal geothermal gradient such that the Moho temperature is ¾600ºC. Steady-state denudation at the Earth’s surface will cause each layer of rock within the crust to cool as it is progressively exhumed by the denudation. Under steady-state conditions at a constant denudation rate U , Stu¨we et al. (1994) derived an equation for the

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temperature T at depth z, which can be written as: Ł T To ð 1 exp. U H =/ D 1 exp. U z=/ T1 To (5) To is the temperature at the Earth’s surface, T1 is the assumed constant temperature at depth H , and  is the thermal diffusivity of the crust which is ¾10 6 m2 s 1 . The temperature profile caused by sedimentation at a steady rate can be analogously calculated with a negative value for U . We model the eastern Greater Caucasus and the surrounding depocentres to the north and south as in Fig. 8. We regard this mountain range as 200 km wide, bounded by depocentres each 100 km wide. We assume a present-day spatial average uplift rate of ¾1 mm a 1 for the eastern Greater Caucasus, and regard their crustal thickening as isostatically compensated such that their spatial average present-day crustal thickening rate is 6 mm a 1 . The required present-day influx of lower crust to sustain this crustal thickening rate is thus 200 km ð 6 mm a 1 or 1200 m3 a 1 per metre of distance parallel to the range front. This flow is assumed to be symmetrical, with an influx Q of 600 m2 a 1 from either flank of the mountain range. Sedimentation rates in the basins adjoining the eastern Greater Caucasus are poorly resolved, not least due to the problems with chronology identified earlier. However, 4 km or more of Kimmerian and younger sediment are known to have accumulated in much of the Kura basin and central Caspian Sea (e.g., Nalivkin, 1973, pp. 627–631; Maslyaev, 1979). We thus adopt a nominal sedimentation rate of 0.6 mm a 1 for both depocentres in the model in Fig. 8. As the surfaces of both depocentres are close to sea level, we presume that a steady state regime exists in both basins where the outflow of lower crust balances the sedimentation. The crustal thickness and hence the Moho temperature thus remain constant. We thus apply Eq. 5 to calculate the depth of the base of the brittle layer, for T D 300ºC, assuming D is the steady crustal thickness of 35 km, and T1 is the Moho temperature of 600ºC. The base of the brittle layer beneath these basins is thus estimated to be at a depth of 20.5 km. The spatial average denudation rate of the eastern Greater Caucasus is poorly resolved by direct evidence: it is somewhere between the ¾10 mm a 1

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Fig. 8. Model for atectonic uplift of the Greater Caucasus. Figure is not to scale, although key dimensions and other parameter values are shown. Material eroded from the eastern Greater Caucasus mountains is transported by rivers to the flanking depocentres in the Kura Basin and central Caspian Sea (1), which also receive sediment from rivers outside the model region, such as the Volga (Fig. 1) (2). Denudation within the eastern Greater Caucasus perturbs the geothermal gradient, causing the base of the brittle upper crust to advect upward relative to the Earth’s surface (3), in accordance with Eq. 5. Sedimentation in the adjoining depocentres likewise causes the base of the brittle layer to advect downward relative to the Earth’s surface (4). As a result of these effects, the pressure at the base of the brittle layer is greater beneath the depocentres than beneath the eastern Greater Caucasus mountains. The resulting horizontal pressure gradient causes lower crust to flow from beneath the depocentres to beneath the eastern Greater Caucasus mountains (5). Bold arrows schematically depict the sense of lower-crustal flow which, in accordance with previous calculations (Westaway, 1998) is concentrated near the base of the crust where the viscosity is least. Parameter values have been chosen to describe a situation where the inflow of lower crust to beneath the eastern Greater Caucasus exceeds the loss of crustal material due to denudation, enabling the crust beneath the eastern Greater Caucasus to thicken. See text for discussion.

rate of incision of some valleys (Rastvorova and Shcherbakova, 1967) and the zero rate of denudation in localities where Sarmatian marine sediments are preserved (Budagov, 1964). As the altitude of the mountain range continues to increase, we presume that the spatial average denudation rate is less than the 1 mm a 1 assumed spatial average uplift rate, and thus adopt a nominal spatial average denudation rate of 0.4 mm a 1 . As the crustal thickness in the eastern Greater Caucasus is increasing, we cannot apply Eq. 5 assuming a constant Moho temperature. We instead use for T1 a constant asthenosphere temperature of 1300ºC, and for H an assumed lithosphere thickness of 140 km comprising an 80 km thick mantle lithosphere layer below the 60 km of crust. Using Eq. 5, we thus estimate the depth of the base of the brittle layer within the eastern Greater Caucasus as 16.3 km.

As is shown schematically in Fig. 8, we assume that the transition from depocentre to mountain range occurs across a distance L of 50 km. Assuming a uniform crustal density ²c of 2700 kg m 3 , and a value of 9.81 m s 2 for g, the acceleration due to gravity, the pressure gradient P across this transition can be estimated as: PD

²c g∆z b L

(6)

where ∆z b is the 4.2 km decrease in depth of the base of the brittle layer between the depocentre and the eastern Greater Caucasus (20.5–16.3 km). The pressure gradient through this transition can thus be estimated as ¾2.2 kPa m 1 . We estimate the typical thickness, W , of the lower crustal channel across this transition as 20 km, Using Eq. 4, with Q D 600 m2 a 1 from each flank

J. Mitchell, R. Westaway / Tectonophysics 304 (1999) 157–186

of the eastern Greater Caucasus and P D 2:2 kPa m 1 , we deduce that the required flow rate can occur in response to the estimated pressure gradient provided e is no greater than ¾7 ð 1019 Pa s. The proposed mechanism is thus feasible if continental crust of normal thickness has the effective viscosity of ¾1019 Pa s predicted by Westaway (1998). We thus conclude from these first-order calculations that the suggested mechanism passes this initial test of feasibility: that uplift of the eastern Greater Caucasus is being sustained by flow of lower crust from beneath adjoining depocentres. It is evident that many simplifying assumptions have been made in these calculations. Among other approximations, we neglect any southward component of lower-crustal flow from beneath the Kura basin, which is presumably responsible for the uplift of the Lesser Caucasus, and neglect the temperature perturbation caused beneath the eastern Greater Caucasus by the inflow of relatively cool crust from surrounding localities, which will affect the local geothermal gradient. Future, more elaborate, calculations will address these points. Much of the material eroded from the eastern Greater Caucasus is subsequently deposited in the adjoining depocentres, partially balancing the transfer of lower crust in the opposite sense. However, our calculations also require a net influx of sediment into the modelled region, as for the model to cause net crustal thickening in the eastern Greater Caucasus the denudation rate of these mountains is required to be less than the sedimentation rate in the adjoining depocentres. We presume that the sediment load transported into the modelled region by large rivers such as the Volga (Fig. 1) and the Kura and Araks (Fig. 3) enables this to be possible. Even with this excess of sedimentation over denudation, the crust beneath the eastern Greater Caucasus will eventually thicken to the point when the pressure at the base of its brittle layer is the same locally as beneath the adjoining depocentres. When this point is reached, the axis of the Greater Caucasus will stop uplifting. However, crust will continue to be driven by the pressure gradient between the depocentres and the flanks of the mountain range, causing continued uplift of these flanks. The width of mountain range at the limiting altitude will thus ultimately broaden. We do not consider it likely that this limiting situ-

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ation has yet been reached, because observed rates of valley incision are greatest along the axis of the Greater Caucasus rather than along the flanks (e.g., Rastvorova and Shcherbakova, 1967; Philip et al., 1989) The start of this atectonic uplift process may relate to changes in environmental conditions in the Late Miocene. From borehole evidence, Hsu¨ and Giavanoli (1979) and others have concluded that the Black Sea basin became desiccated at the time of the Messinian drawdown of sea-level in the Mediterranean during latest Miocene time. The Meiotian drawdown in level in the Caspian basin, discussed by Nalivkin (1973, pp. 630–632) may mark the same event: as was noted above, in the Soviet chronology, the Meiotian is latest Miocene and thus synchronous with the Messinian, whereas in the Steininger and Papp (1979) chronology the Meiotian stage of the Paratethys chronology predates the Messinian stage of the Mediterranean by several million years. Nonetheless, if the water levels in the Black Sea and Caspian Sea basins were briefly drawn down at the same stage of Late Miocene time, the resulting increase in subaerial relief would be expected to accompany an increase in the denudation rate of the proto-Greater Caucasus. This set of circumstances may have been sufficient to initiate the cycle of coupled sedimentation and denudation depicted in Fig. 8. However, before this possibility can be tested, it will be necessary to reconcile the conflicting time scales for the Late Miocene. Finally, we investigate the relationship between the timings of uplift and magmatism in the Greater and Lesser Caucasus. From earlier discussion, in the Lesser Caucasus the magmatism began at or shortly after 1.5 Ma, and became most intense around 0.8 Ma; in the Greater Caucasus it began before 3 Ma and became most intense around 2.8 Ma. Using Fig. 8, we estimate that ¾2.0–2.5 km of uplift has occurred in the Greater Caucasus since the start of magmatism. At the start of magmatism, the altitude of the Greater Caucasus was thus ¾1.7–2.2 km. In the Lesser Caucasus, the present-day surface altitude excluding volcanic edifices would be ¾2.0–2.5 km. Data in Table 5 indicate that this altitude has increased by ¾0.5 km since Late Akchagyl time, around the start of volcanism (Fig. 6). The surface altitude at the start of volcanism was thus ¾1.5–2.0

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km, roughly the same as is estimated for the Greater Caucasus. These estimates of the surface altitude at the start of volcanism indicate a crustal thickness of ¾45 km. Atectonic thickening of the continental crust keeping the mantle lithosphere thickness constant will raise the temperature at the Moho and in each layer of mantle lithosphere. We thus conclude that previous suggestions (e.g., Koronovskiy and Demina, 1996), that magmatism in the Caucasus has resulted from crustal thickening, are indeed correct. A temperature increase in the uppermost mantle, consistent with an increase in crustal thickness to from ¾35 to ¾45 km, seems to have been required to initiate this process in this region.

8. Discussion This study proposes a new mechanism for mountain building: as a result of flow of lower crust to beneath a mountain range from beneath its surroundings, in response to a horizontal pressure gradient at the base of the brittle layer oriented in this sense. This pressure gradient is established by the effect of erosion perturbing the geothermal gradient so as to reduce the depth of the base of the brittle layer beneath the mountain range, and by a corresponding increase in depth of the base of the brittle layer beneath surrounding regions. Earlier calculation established that for this process to be feasible on the scale of the Caucasus, a weak lower crust is required as well as high rates of erosion and sedimentation. It is important to contrast this deduction with earlier views. The earliest detailed quantitative analyses of continental deformation used the ‘thin viscous sheet’ model (e.g., England and McKenzie, 1982), in which the continental lithosphere was modelled as a viscous layer. Deformation of this layer at each time step of the model was assumed to depend on the relative motion of its edges, due to plate motions, and on the body forces which develop due to crustal thickness variations caused by earlier deformation. One outcome from this work was the deduction (e.g., by England et al., 1985; and Houseman and England, 1986) that plate convergence will initially cause crustal thickening within the plate boundary zone, but once a limiting surface altitude is reached, the body forces due to the topography will cause the

area of thickened crust to widen perpendicular to the plate boundary. By making appropriate approximations, Gratton (1989) derived analytic equations for the sense of deformation, as a function of the convergence rate V , required to produce an isostatically-compensated saw-tooth mountain range of height H and width A. Gratton (1989) assumed that such topography is supported by dynamic equilibrium between the viscous stress caused by crust flowing into this deforming zone due to the convergence of its boundaries, and the stress caused by the horizontal pressure gradient established within the crust by the topography (which on its own would cause crust to flow out of the deforming zone). Both stress terms were spatially averaged, and the crust was assumed to have a temperature-independent rheology with no upper-crustal brittle layer. For a model incorporating vertical gradients in the horizontal velocity within an isoviscous crust (with viscosity ), Gratton (1989) showed that the resulting scaling behaviour of the topography satisfies: 1=4  DV 2 ½3 t (7) HD 2.1 C ½/²g and:  AD

32D 3 V 2 ½²gt 3 .1 C ½/3 

1=4 (8)

where ² and D are the density and initial thickness of the crust, and ½ equals .²m ²/=², where ²m is the density of the mantle. Gratton (1989) considered that a realistic value for  was ¾1021 Pa s. Taking D D 30 km, ² D 2700 kg m 3 , g D 10 m s 2 , and ½ D 0:2, Eq. 7 indicates that a convergence rate V D 6 mm a 1 would create ¾4.3 km of topography in ten million years, reflecting the overall uplift history of the Greater Caucasus. Furthermore, from Eq. 8, the resulting mountain range is estimated to have developed a width of ¾230 km by this time, also comparable to the Greater Caucasus. However, it is evident that Eq. 7 predicts that H / t 1=4 , or @ H =@t / t 3=4 . A model of this form is thus unable to reproduce the observed progressive increase in the uplift rate of the Greater Caucasus (Fig. 7). Gratton (1989) also showed that ‘thin viscous sheet’-type models predicted similar scaling behaviour to models in which crustal flow creates

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vertical gradients in horizontal velocity, and that the form of the topography produced by assuming a linear rheology does not differ much from what is produced by assuming power-laws between stress and strain rate. The earlier result (e.g., Houseman and England, 1986), that progressive broadening of zones of surface uplift following continued convergence requires ‘thin viscous sheet’-type models with power-law rheologies, was thus refuted. The model of Gratton (1989) did not include an upper-crustal brittle layer. Subsequent studies (e.g., Westaway, 1994a,b, 1995, 1998; and this study) have indeed highlighted the importance of lateral variations in the thickness of this layer in establishing horizontal pressure gradients to drive lower-crustal flow. Because the model of Gratton (1989) did not consider erosion, if a brittle layer were added to this model, it would be thicker over the high topography than in its surroundings due to the reduced geothermal gradient. This would set up a pressure gradient which would drive lower crust out from beneath the high topography, which would cause some local relative crustal thinning and some crustal thickening in the surroundings. The result would be the creation of more subdued relief than the model of Gratton (1989) predicts. Some studies (e.g., Zhao and Morgan, 1987; Westaway, 1995) have indeed used this feature of lower-crustal flow to derive upper bounds to the viscosity of the lower crust from lateral variations in surface altitude in elevated regions such as the Tibetan plateau. Bird (1991) made the next significant contribution to analysing lower-crustal flow re´gimes linking regions of thickened crust and their surroundings. He assumed that lateral variations in topography had been previously created by some earlier process, and examined the subsequent evolution of this topography due to lower-crustal flow. His finite-element modelling showed lower crust flowing from regions of thickened crust to regions where the crust was initially thinner, thus causing the initial crustal thickness variations to decay. Bird (1991) thought that this effect was due to the lateral gradient in pressure at the Moho, which is deeper beneath the thicker crust and thus locally under higher pressure. However, it should be clear that flow in a confined fluid depends on lateral pressure gradients at its top, not its base. The true physical reason why the model of

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Bird (1991) works appears to be that it incorporated lateral variations in the initial geothermal gradient, which varied inversely with the initial crustal thickness. The initial pressure at the base of the brittle layer was thus greatest in regions with the thickest crust, thus determining the flow sense. Westaway (1994a) appears to have been the first study to point out that loading effects caused by sedimentation, or negative loads caused by erosion, may create lateral variations in pressure at the base of the brittle layer which can outweigh the effect of crustal thickness variations, thus causing lower crust to flow to beneath regions with relatively thick crust. Royden (1996) devised a detailed analytic model for predicting lower-crustal flow regimes associated with continental convergence. Although she treated some aspects of this problem to high precision, erosion of topography was disregarded; the brittle upper crust was modelled as a high-viscosity fluid, and the temperature-dependence of lower-crustal viscosity was represented approximately assuming an exponential decay of viscosity with depth. Royden (1996) showed that if the lower crust has a high viscosity, sawtooth mountains develop as a result of convergence, but if it has a low viscosity, then high topography spreads laterally producing a plateau of roughly uniform altitude. These results thus confirmed the earlier conclusions of Gratton (1989) based on much simpler considerations. Royden et al. (1997) used this model to suggest that the viscosity of the lower crust beneath Tibet is so low that the brittle upper crust can move independently of the underlying mantle lithosphere, thus explaining the eastward extension of Tibet and the accompanying crustal deformation around this plateau’s eastern margin: another conclusion which had already been established, on the basis of simpler considerations, by Westaway (1995). The conclusion of Royden (1996), that low topographic gradients imply that the lower crust is weak, and high topographic gradients imply that it is strong, is contradicted by the present study: for lower-crustal flow induced by surface processes, the lateral variations in relief will reflect the lateral variations in surface processes, and can thus be abrupt even if the lower crust is weak. To the best of our knowledge, Avouac and Burov (1996) were the first to assess the relative importance of convergence and erosion in determining the

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evolution of mountain ranges. They developed an elaborate coupled numerical model linking crustal rheology, crustal deformation, erosion from a mountain range and sedimentation in its surroundings. They showed that with no erosion, if the lower crust is weak it will flow outward from the axis of the high topography (as was demonstrated by Bird, 1991), causing gradients in topography and crustal thickness to decay with time. Second, with erosion occurring as well as convergence, the different isostatic response leads to net inflow of lower crust to beneath the high topography, causing progressive surface uplift. However, third, they deduced that in the presence of erosion but with no convergence, lower crust will flow out from beneath the high topography, causing the gradual decay of this topography. The result of Avouac and Burov (1996) that with a zero convergence rate, lower-crustal flow will cause topography to decay, differs fundamentally from our result that the relief of the Caucasus is developing with a zero convergence rate. There are two main reasons for this difference. First, Avouac and Burov (1996) stated that one of the approximations made by their model is to disregard perturbations to the geothermal gradient caused by surface processes. In their model, the thickness of the brittle layer increases with crustal thickness, and (in the absence of convergence) the resulting lateral pressure gradient will inevitably drive lower crust out from beneath a mountain range. Avouac and Burov (1996) neglected perturbations to the geothermal gradient caused by surface processes because they thought that these would only become significant on time scales longer than the ¾10 Ma durations of their model runs. We have assumed instead that the geothermal gradient profile in the crust adjusts continually to reflect the changing rates of surface processes, and at all times is equal to that which would exist if the contemporaneous rate of erosion and sedimentation had persisted indefinitely. Calculations by one of us (R.W.), which will be published elsewhere, indicate that the lower-crustal flow response to changing rates of surface processes can indeed be significant on time scales of <1 Ma. The second major reason is that Avouac and Burov (1996) assumed a nonlinear, power-law, rheology for the lower crust, whereas we have assumed (following Westaway, 1998) a linear rheology. The justification for using a linear rheol-

ogy for lower-crustal flow was given by Westaway (1998). It is partly based on general considerations of first principles, and partly from a re-examination of experimental rock mechanics data; and is not repeated here. As Westaway (1998) demonstrated, the method for calibrating non-linear rheologies leads, for deformation at geological strain rates, to overestimation of the viscosity at the depths within the lower crust where flow is concentrated by several orders of magnitude. For instance, in fig. 7 of Avouac and Burov (1996) a lateral variation in brittle-layer thickness of ¾4 km in ¾100 km distance induces lower-crustal flow which, in the part of the model with this pressure gradient, has a peak velocity of ¾1 mm a 1 and mean velocity of ¾0.3 mm a 1 in lower crust with a mean thickness of ¾18 km. From Eq. 6, the horizontal pressure gradient in this model is ¾1.1 kPa m 1 , and thus, from Eq. 4, the effective viscosity of its lower crust can be estimated as ¾3 ð 1021 Pa s. If the lower crust had such a high effective viscosity, the lateral pressure gradient induced by surface processes would not be able to drive lower-crustal flow fast enough to sustain the observed crustal thickening rate in the Greater Caucasus. The combination of using the correct sense of lateral pressure gradient together with a low effective viscosity for the lower crust makes our scheme in Fig. 8 mechanically feasible.

9. Conclusions We have determined K–Ar ages for Pleistocene volcanic rocks at three sites in Armenia (Table 3). Our dates suggest that the local Quaternary volcanism began before 1.1 Ma and reached its peak around 0.8 Ma. They also provide absolute age control for other features in and around this study region, which have not previously been reliably dated, by using correlation schemes between volcanism, glaciations, and the stratigraphy of the Caspian basin, which have previously established the relative chronology of events. We suggest a ¾1.2 Ma timing for the Late Akchagyl transgression of the Caspian Sea. We tentatively correlate this with the melting event following glaciation during stage 36 of the oxygen isotope timescale, which was thus the first time during the Pleistocene when eastern Europe was covered by a

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lowland ice sheet. Time-averaged since ¾1 Ma, the eastern Greater Caucasus mountains are thus shown to have uplifted at ¾0.6 mm a 1 and the Lesser Caucasus at ¾0.3 mm a 1 . We suggest that this uplift has no direct relationship to plate convergence in the region, but is instead being caused by inflow of lower continental crust to beneath these uplifting regions from beneath adjoining depocentres.

Acknowledgements We thank Rob Ridley for his assistance in K–Ar dating. Supported by Natural Environment Research Council grant GR3=7345.

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